The Moho in Extensional Tectonic Settings: Insights from Thermo-Mechanical Models

TECTO-125938; No of Pages 47 Tectonophysics xxx (2013) xxx–xxx

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Review Article The Moho in extensional tectonic settings: Insights from thermo-mechanical models

Sierd Cloetingh a,⁎, Evgenii Burov b,c, Liviu Matenco a, Fred Beekman a, François Roure a,e, Peter A. Ziegler d,† a Netherlands Research Centre for Integrated Solid Earth Sciences, Faculty of Geosciences, Utrecht University, Budapestlaan 4, 3584 CD Utrecht, The Netherlands b UPMC Univ Paris 06, ISTEP UMR 7193, Université Pierre et Marie Curie, F-75005 Paris, France c CNRS, ISTEP, UMR 7193, F-75005 Paris, France d Kirchweg 41, 4102 Binningen, Switzerland e IFPEN, 1-4 Avenue de Bois-Préau, 92852 Rueil-Malmaison Cedex, France article info abstract

Article history: The lithospheric memory is key for the interplay of lithospheric stresses and rheological structure of the Received 22 November 2012 extending lithosphere and for its later tectonic reactivation. Other important factors are the temporal and Received in revised form 12 June 2013 spatial migration of extension and the interplay of rifting and surface processes. The mode of extension Accepted 13 June 2013 and the duration of the rifting phase required to lead to continental break-up are to a large extent controlled Available online xxxx by the interaction of the extending plate with slab dynamics. The ﬁnite strength of the lithosphere has an important effect on the formation of extensional basins. This applies both to the geometry of the basin shape Keywords: Lithosphere rheology as well as to the record of vertical motions during and after rifting. We demonstrate a strong connection between Lower crustal ﬂow the bulk rheological properties of Europe's lithosphere and the evolution of some of Europe's main rifts and Moho imagery back-arc systems. The thermo-mechanical structure of the lithosphere has a major impact on continental Mode of rifting break-up and associated basin migration processes, with direct relationships between rift duration and extension Stress regimes velocities, thermal evolution, and the role of mantle plumes. Compressional reactivation has important conse- Lithospheric memory quences for post-rift inversion, borderland uplift, and denudation, as illustrated by poly-phase deformation of extensional back-arc basins in the Black Sea and the Pannonian Basin region. © 2013 Elsevier B.V. All rights reserved.

Contents

1. Introduction ...... 0 2. Constraints by deep seismic imaging, potential field data and mantle tomography on the architecture of the Moho, volcanic underplating and the occurrence of residual lithospheric slabs in rift basins, passive margins and intracontinental sags ...... 0 2.1. Architecture of the Moho and lower crust in extensional settings, as revealed by deep seismic reflection profiles...... 0 2.2. Evidence for volcanic underplating from refraction data and inversion of potential field data ...... 0 2.3. Long lasting subsidence of intracratonic sags induced by residual lithospheric slabs ...... 0 2.4. Negative inversion of former thrust belts ...... 0 3. Rifting and extensional basin formation and its role in the evolution of the continental lithosphere ...... 0 4. Extensional tectonics: concepts and global-scale observations ...... 0 4.1. Extensional basin systems ...... 0 4.1.1. Thermal thinning and stretching of the lithosphere: concepts and models ...... 0 4.1.2. Syn-rift subsidence and duration of rifting stage ...... 0 4.1.3. Post-rift subsidence ...... 0 4.1.4. Post-rift compressional reactivation of extensional basins ...... 0 4.1.5. Finite strength of the lithosphere during extensional basin formation ...... 0 4.1.6. Rift-shoulder development and architecture of basin fill...... 0 5. Rheological stratification of the lithosphere and basin evolution ...... 0 5.1. Lithosphere strength and deformation mode ...... 0 5.2. Mechanical controls on the evolution of rifts: Europe's continental lithosphere ...... 0

⁎ Corresponding author. Tel.: +31 30 2537314; fax: +31 30 2535030. E-mail address: [email protected] (S. Cloetingh). † Deceased July 19th, 2013.

5.3. Lithospheric folding: an important mode of compressional reactivation of rifts ...... 0 6. Models for continental breakup and rift basins ...... 0 6.1. Extensional basin migration: observations and thermo-mechanical models ...... 0 6.2. Fast rifting and plate separation ...... 0 6.3. Thermo-mechanical evolution and tectonic subsidence during slow extension ...... 0 6.4. Interplay of lower crustal ﬂow and surface erosion in rifts ...... 0 6.5. Breakup processes: timing and mantle plumes ...... 0 6.6. The evolution of the Moho geometry during rifting ...... 0 6.7. Post-rift inversion, intraplate stresses, borderland uplift, and denudation ...... 0 7. Back-arc basin formation and evolution ...... 0 7.1. Black Sea ...... 0 7.1.1. Rheology and back-arc rift basin formation ...... 0 7.1.2. Strength evolution and neotectonic reactivation at the basin margins during the post-rift phase ...... 0 7.2. Modes of basin (de)formation, lithospheric strength, and vertical motions in continental back-arc rifts ...... 0 7.2.1. Neogene development and evolution of the Pannonian Basin ...... 0 7.2.2. Dynamic models of basin formation ...... 0 7.2.3. Stretching models and subsidence analysis ...... 0 7.2.4. Lithospheric strength in the Pannonian–Carpathian system ...... 0 7.2.5. Deformation of the Pannonian–Carpathian system ...... 0 8. Conclusions ...... 0 Acknowledgements ...... 0 References ...... 0

1. Introduction the thermo-mechanical aspects of extensional sedimentary basin development in the context of large-scale lithosphere models for the The Moho, originally defined as an acoustic contrast marking the underlying lithosphere and on tectonic controls on the post-rift evo- compositional boundary between the crust and the underlying mantle, lution of extensional basins. is also of crucial importance for the mechanical state of the lithosphere. We highlight the connection between the bulk rheological proper- Actually, the effect of the inherited crustal fabric and strength appear to ties of the lithosphere and the evolution of extensional basins and be of paramount influence on the nature of tectonic deformation in the find that the finite strength of the lithosphere plays an important plate interiors. A particularly important parameter for lithospheric role in their development. This applies both to the geometry of the strength is the ratio of the thickness-dependent crustal and lithospheric basin shape as well as to the record of vertical motions during and mantle strength (e.g. Burov, 2011)anditisexactlythisratiothatis after lithospheric extension. strongly affected by crustal thinning during the rifting process. In addi- Below we give a brief review of concepts for the rheological struction, changes in thermal regime characteristic for rifting processes leave ture of continental lithosphere and discuss insight gained from global a strong imprint on the temporal and spatial variations in crustal and models of strength distribution, highlighting the impact rifting pro- lithospheric strength during and after the rifting phases. cesses have on the lithospheric strength distribution. This is followed The notion that rifting is a very important mechanism for sedimen- by a discussion of constraints on the duration of rifting based on case tary basin formation with key implications for the basin fill and thermal studies (e.g. Ziegler and Cloetingh, 2004), the effects of slow versus regime is of immediate importance for the assessment of the hydrocar- fast extension on lithospheric strength and crustal configuration bon habitat. This has boosted rifting studies in the context of global (van Wijk and Cloetingh, 2002; Van Wijk et al., 2004), and an analysis exploration for hydrocarbons. Deep seismic profiling, often carried out of the interaction between surface erosion and lower crustal flow in in close collaboration between academia and industry (e.g. Thybo and extensional regimes (see also Burov and Cloetingh, 1997; Burov and Nielsen, 2009; Thybo et al., 2000), has been an essential component Poliakov, 2001). We then analyze the Pannonian basin and the Black of these efforts. As a result, many data are now available to provide Sea back-arc basins, integrating observational constraints on crustal constraints on the timing and duration of the rifting phases and overall structure and fabric provided by deep seismic profiles with inferences geometry of rift basins. from thermo-mechanical models. At the same time, advances in seismic tomography studies of the mantle structure at convergent plate boundaries, combined with case 2. Constraints by deep seismic imaging, potential field data and history analyses, have revealed a much stronger coupling between the mantle tomography on the architecture of the Moho, volcanic subducting lower lithospheric plate and the overriding upper litho- underplating and the occurrence of residual lithospheric slabs in spheric plate that affects a much larger area of the latter than originally rift basins, passive margins and intracontinental sags envisaged. The thermo-mechanical structure of the lithosphere is of fundamental importance for the interaction of both upper mantle and The localization, initiation, development and current architecture lithospheric processes (Burov and Cloetingh, 2009, 2010; Burov et al., of sedimentary basins and orogens directly rely on the thermo- 2007) as well as for the assessment of the horizontal tectonic stress im- mechanical behaviour of the continental lithosphere, which is in turn pact on lithosphere deformation (Burov and Cloetingh, 2009; Cloetingh depending on the local thermal regime (i.e. lithosphere thickness or and Burov, 2011). depth to the 1300 °C isotherm; Artemieva et al., 2006), and on the ver- This paper reviews recent advances in modelling of the initiation tical and lateral distributions of the various lithologies encountered in and evolution of extensional basins in a lithospheric context. We the lower crust, crystalline basement and overlying sedimentary rocks first summarize key features of the architecture of rifted basins and (Allemand and Brun, 1991; Bertotti et al., 2000; Burov and Watts, examine the record of vertical motions during and after rifting in 2006; Cloetingh and Banda, 1992; Davy and Cobbold, 1991; Kusznir the context of stretching models developed to quantify the develop- and Karner, 1985; Huismans and Beaumont, 2003, 2008; Huismans ment of extensional basins. This will be followed by a discussion on et al., 2005; Ranalli, 1995; Ranalli and Murphy, 1987).

At lithospheric scale, the main decoupling horizons operating at COT (Continental–Oceanic Transition) in the inner part of these plate boundaries but also within lithospheric plates mostly comprise orogens. the Moho surface and the ductile lower crust, as well as pre-existing Ultimately, COCORP profiles in the Basin and Range province tectonic fabrics still preserved in the brittle crust, such as older (Allmendinger et al., 1983, 1987; Brown et al., 1986; Hauser et al., low-angle extensional detachments or thrust systems which are likely 1987) and combined field studies and seismic imagery in the Apennines to be positively or negatively inverted during subsequent compression- (Scrocca et al., 2005; and references therein) and Alboran, Algerian al (Brooks et al., 1988; Colletta et al., 1997; Cooper and Williams, 1989; basins and Tyrrhenian basins (Déverchère et al., 2013; Medaouri et al., Coward, 1983; Gillchrist et al., 1987; Letouzey, 1990; Letouzey et al., in press; Ranero et al., 2013; Roure et al., 2009) have also contributed 1990; Roure and Colletta, 1996; Van Wees, 1994; Vially et al., 1994)or to better understand the change operating from compression to exten- extensional episodes, respectively (Burg et al., 1994; Dewey, 1988; sion in over-thickened tectonic wedges and during the opening of McClay et al., 1986; Menard and Molnar, 1988; Platt and Vissers, back-arc basins. 1989; Seguret et al., 1989; Séranne et al., 1989; van den Driessche and Brun, 1991). 2.2. Evidence for volcanic underplating from refraction data and inversion of potential field data

2.1. Architecture of the Moho and lower crust in extensional settings, as Rapid vertical and lateral changes can occur in the lithologies and revealed by deep seismic reflection profiles seismic velocities within the upper and the lower crust, both beneath rift basins and passive margins. For instance, the consolidation of volcanic In contrast to Moho depth predictions based on the inversion of melts at or near the Moho interface can result in anomalously dense potential data, deep seismic reflection profiles provide a direct image bodies at the base of the crust, likely to induce long-term subsidence, of the architecture of the Moho and lower crust beneath many rift independently to temporary effects of lithospheric cooling and post-rift basins and passive margins. Furthermore, refraction and tomographic thermal subsidence. Seemingly, hydration/serpentinisation of denudated data help also to constrain robust velocity models that can be used to mantle at the Continent–Ocean Transition can lead to anomalously slow convert time sections into depth sections, and to better control the velocities for the mantle, which should not be misinterpreted as radio- current attitude of the Moho, as well as the local occurrence of anoma- genic lower crust, which is instead always characterized by seismic veloc- lous high velocity/dense bodies that account for granulites in the lower ities higher than the brittle upper crust. crust, underplated volcanics that have been consolidated in the vicinity Time sections beneath the Parentis basin (Bois and Gariel, 1994; of the Moho interface, or even remnants of eclogitized lithospheric slabs Fig. 1A) and the Dienpr–Donets (Stephenson et al., 2001; Fig. 1B) that are still attached beneath the Moho. basin display a flat Moho, which would instead become shallower Despite the fact that the initial stage of the lithosphere prior to the when the seismic profile is converted into a depth section in the onsetofriftingmaybedifficult to assess, at least the present day atti- case of the Parentis basin, due to the slow velocity of overlying Meso- tude of the Moho and lithosphere thicknesses documented by geophys- zoic sediments. In the case of the Dniepr–Donets basin. It is not clear, ical tools constitute unique data to test/validate thermo-mechanical however, whether the amount of underplating and salt (high velocity models of lithosphere/crustal evolution in extensional systems. For rocks) is sufficient to balance the effect of slower velocity clastic and instance, in the framework of the French Ecors program deep seismic carbonate sediments, and therefore, whether the Moho surface is still profiles have been recorded in the eighties in extensional systems, ultimately flat, or would instead become uplifted or down-flexed in two different segments of the West European Oligocene rift system once properly depth converted. across the Rhine (Brun et al., 1991, 1992) and Bresse grabens Therefore, it is of prime importance to define a robust velocity model (Bergerat et al., 1989, 1990), in the Bay of Biscaye across the Albian at crustal scale when converting time images into depth images. In depocenter of the Parentis Basin (Bois and Gariel, 1994), and in the addition to refraction data that are frequently recorded along the Gulf of Lions across the northern passive margin of the West Mediterra- same regional profiles as deep reflection seismic, the joint inversion nean Basin (Burrus, 1989; De Voogd et al., 1991; Séranne, 1999). of gravimetric and magnetic data is also frequently used to trace the Also, the rapid development of petroleum exploration in the deep distribution of dense bodies in the vicinity of the Moho and improve offshore by the industry has significantly promoted the acquisition of our overall understanding of the thermal state of the crust and over- crustal-scale geophysical data on various segments of the conjugate lying sedimentary basins, as for instance in the South Atlantic and passive margins of the Atlantic between Europe and North America Norwegian margins (Maystrenko et al., in press; Scheck-Wenderoth (Etheridge et al., 1989; Faleide et al., 2008; Péron-Pinvidic and and Maystrenko, 2008) or the Asia continent (Stolk et al., 2013). Manatschal, 2009; and references therein), and between Africa and South America (Castro, 1987; Franke et al., 2012; Moulin et al., 2010; 2.3. Long lasting subsidence of intracratonic sags induced by residual Ueipas-Mohriak et al., 1995; Unternehr et al., 2010, and references lithospheric slabs therein), as well as across the conjugate passive margins of Australia and Antarctica (Direen et al., 2011; Espurt et al., 2012, and references Many passive margins (e.g. the South Atlantic and South Australia therein), thus providing a clear picture of the current architecture of margins) are characterized by a long lasting sag episode, following mature rifted margin systems with either volcanic-rich or volcanic- major stretching and associated crustal tilted fault block formation. poor signatures. Despite the fact that their petroleum potential is Although the controls of sag development are still debated, two obvious more limited, young oceanic basins like the Gulf of California or the explanations can be proposed, depending on the assumptions made on Gulf of Aden have also been studied recently by means of crustal-scale the velocity models and the actual architecture of the Moho beneath geophysical surveys (Leroy et al., 2012, and references therein). the sags. The first explanation assumes that either the Moho remains Deep oceanic drilling off the Galicia margin (Boillot et al., 1988; flat beneath the sag basin or the dense anomalies relate to volca- Manatschal and Bernoulli, 1999; and references therein) and field stud- nic underplating, as suggested by Maystrenko et al. (in press) and ies in the Alps and the Pyrenees (Lagabrielle and Bodinier, 2008; Scheck-Wenderoth and Maystrenko (2008) accounting for a buried Manatschal et al., 2001; Manatschal, 2004; Manatschal et al., 2011) load beneath the sag. Alternatively, it is assumed that the Moho is have also contributed to control the lithologies and tectonic fabrics of uplifted beneath the sag, with the sedimentary loading in the sag lower crustal and upper mantle rocks which are still currently exposed resulting in a lateral flow of the lower crust away from the depocenters. at the continent–ocean transition in the deep offshore, or are instead Last but not least, new tomographic models beneath the Paris Basin preserved as allochthonous tectonic units still preserving the former have documented a strong velocity anomaly in the upper mantle, which

Fig. 1. A) Line drawing and simplified geological interpretation of the Ecors profile crossing the Parentis Basin. Notice the relatively flatMohoonthetimesection,andtheoverallasymmetryof the basin. Due to crustal stretching, Albian sediments rest almost directly on top of the lower crust and continental mantle lithosphere (Bois and Gariel, 1994). B) Line drawing and simplified geological interpretation of the DOBRE deep seismic profile crossing the Dniepr–Donets graben. Notice the apparent flat Moho beneath the main Devonian depocenter on the time section. High velocity salt occurs in the sedimentary infill of this basin, whereas a dense high velocity body interpreted as underplated volcanics has been identified. Ultimately, post-Devonian positive inversion of this basin has also been identified, accounting also for reverse faulting (Maystrenko et al., 2003; Stephenson et al., 2001; Stovba and Stephenson, 2003). is best interpreted as a remnant of the former Hercynian lithospheric substantial progress was made in the understanding of thermo- slab. The preservation of such a dense sub-crustal load since the Carbon- mechanical processes controlling the evolution of rifts and extensional iferous is probably the mechanism accounting for the Mesozoic subsi- sedimentary basins and the isostatic response of the lithosphere to dence of the overlying basin, where no Triassic nor Jurassic normal surface loads such as sedimentary basins (Watts, 2001). Much of this faults have ever been documented, and where Alpine foreland litho- progress stems from improved insight into the mechanical properties spheric buckling can only be advocated to explain the Cenozoic subsi- of the lithosphere, from the development of new modelling techniques, dence pattern of the basin (Averbuch and Piromallo, 2012; Fig. 2A; and from the evaluation of new, high-quality datasets discussed above Bourgeois et al., 2007; Fig. 2B). from previously inaccessible areas of the globe. Below we focus on tectonic models of processes controlling the evolution of extensional 2.4. Negative inversion of former thrust belts basins. After the realization that subsidence patterns of Atlantic-type mar- Negative inversion of Paleozoic thrusts inherited from the Caledonian gins, corrected for effects of sediment loading and paleo-bathymetry, and Hercynian orogenies has been evidenced in the Celtic Sea by BIRPS, displayed the typical time-dependent decay characteristic of ocean- SWAT and WAM deep seismic profiles (Bois et al., 1991; Hobbs and floor cooling (Sleep, 1971), a large number of studies were undertaken Klemperer, 1991; Lefort et al., 1991; Fig. 3), in association with a lateral aimed at restoring the quantitative subsidence history of rifted basins flow of the lower crust and development of a Mesozoic sag basin in the on the basis of well data and outcropping sedimentary sections. With British Channel. Surface observations and wells correlations have also pro- the introduction of backstripping algorithms (Bond and Kominz, vided evidence for the poly-phase reactivation of the Hercynian front (faille 1984; Steckler and Watts, 1978) in the late 1970s' and early 1980s', a du Midi), both negatively during Mesozoic extensional episodes and posi- phase of basin analysis commenced that aimed at backward modelling tively during Paleogene compressional episodes (Averbuch et al., 2004). by reconstructing the tectonic basin subsidence from sedimentary sequences. These quantitative subsidence histories provided con- 3. Rifting and extensional basin formation and its role in the straints for the development of conceptually driven forward basin evolution of the continental lithosphere models. For extensional basins this commenced in the late 1970s, with the appreciation of the importance of lithospheric thinning Rifting and extensional basin formation is a key factor in the geolog- and stretching to basin subsidence. After the initial mathematical ical evolution of the continental lithosphere. During the last decades, formulation of stretching concepts in forward modelling of extensional

Fig. 2. A) Sketch diagram outlining the occurrence of a high velocity dense body in the upper mantle beneath the Paris Basin, as evidenced by mantle tomography, which is interpreted as a relic of the former Hercynian slab (after Averbuch and Piromallo, 2012). B) Top: Patterns of lithospheric folding and orientation of folding axis in the northwest European Alpine foreland. Bottom: cross-section A–A′ (see top ﬁgure for location), from the Alpine thrust front to the Paris Basin, outlining the large wave-length Cenozoic buckling of the European foreland lithosphere (after Bourgeois et al., 2007).

Fig. 3. Line drawing from SWAT and WAM deep seismic profiles across the Celtic Sea and British Channel, outlining the negative inversion of Paleozoic thrusts inherited from the Caledonian and Hercynian orogenies (after Bois et al., 1991; Hobbs and Klemperer, 1991; Lefort et al., 1991). Notice also the lateral flow of the lower crust and coeval development of a Mesozoic sag basin. basins (McKenzie, 1978), basin fill simulation focused on the interplay seismics, wells, outcrops); this required a close cooperation between between thermal subsidence, sediment loading, and eustatic sea-level academic and industrial research groups (see Cloetingh and Ziegler, changes. To simulate episodic and irregular anomalies commonly 2007; Cloetingh et al., 2010; Roure, 2008; Roure et al., 2010). observed in otherwise smooth post-rift subsidence curves, changes in Regarding thermo-mechanical controls on continental breakup and sediment supply and eustatic sea-level fluctuations were implemented associated basin migration processes, we summarize the results of some in an effort to identify their potential tectonic origin (intraplate stress- of our studies on the continental margin of Norway, one of the best doc- induced basin subsidence/uplift). In another approach a subsidence umented in the world owing to concerted efforts by academia and in- curve was used as input, with the basin-modelling package essentially dustry in the context of hydrocarbon exploration and production. This filling the adopted accommodation space (e.g., Lawrence et al., 1990). rifted margin provides close constraints on the timing and magnitude For the evolution of extensional basins this approach made a clear dis- of extension, extension velocities and its thermal evolution, as well as tinction between their syn-rift and post-rift stages, relating exponen- on its post-rift compressional reactivation and partial inversion that tially decreasing post-rift tectonic subsidence rates to a combination had repercussions on the uplift and erosion of its borderland. of thermal equilibration of the lithosphere–asthenosphere system and As any other passive margin, however, it is difficult in Norway and to lithospheric flexure (Watts et al., 1982). its Greenland conjugate margin to constrain the initial architecture of Similar simplifications were adopted to describe the syn-rift phase. the crust and lithosphere prior to the onset of rifting. That would be a In the initial version of the stretching model (McKenzie, 1978), litho- pre-requisite to provide consistent input of thermal or analogue model- spheric thinning was described as resulting from instantaneous exten- ling. For instance, a prime question is the actual control of Caledonian sion of a semi-infinite half-space with infinite Péclet number. Later thrusts (Anderson et al., 1992) in the localization of the rift-related more complex tilted fault block models were introduced that decoupled normal faults and the extent of upper Paleozoic series in the offshore the crust from the lithospheric mantle during the rifting stage (e.g. Brun domain. Also, the timing of the unroofing of the Paleozoic series and and Nalpas, 1996; Burov and Cloetingh, 1997; Kaufman and Royden, crystalline basement onshore Norway and Sweden is still debated. The 1994; Kusznir et al., 1991). These rather theoretical models did, how- other major question being related to the initial extent and thickness ever, not take into account the history of rifted basins and obviously of the Jurassic series onshore, prior to the exhumation of their underly- lacked a number of important components of lithosphere mechanics. ing basement as due to the more recent uplift. A notable feature of most modelling approaches was their emphasis Poly-phase deformation of extensional basins is discussed further on the basin subsidence record and their very limited capability to han- on in this paper with examples of the Black Sea and Pannonian– dle differential subsidence and uplift patterns in a process-oriented, Transylvanian back-arc basins. An intriguing feature of these basin sys- internally consistent manner (e.g., Kusznir and Ziegler, 1992). Subse- tems is the asymmetric relationship between the syn-rift and post-rift quent modelling focused on the quantification and coupling of subsidence patterns that reflect overall thermo-mechanical asymmetry. lithosphere processes and their near-surface expression in response to This demands an assessment of the rheological controls on basin forma- tectonic controls on basin fill. This demanded a process-oriented ap- tion and of implications for the large-scale basin stratigraphy and rift proach, linking different spatial and temporal scales in the basin record. shoulder dynamics. In this context the role of intraplate stresses and In this it was crucial to test and validate model predictions against the strength evolution of the lithosphere during the post-rift phase high-quality data providing information at deeper crustal levels (deep are of particular interest for the understanding of the neotectonic reflection- and refraction-seismics) and the basin fill (reflection- reactivation of the Black Sea Basin system. An overview of the interplay

Please cite this article as: Cloetingh, S., et al., The Moho in extensional tectonic settings: Insights from thermo-mechanical models, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.06.010 S. Cloetingh et al. / Tectonophysics xxx (2013) xxx–xxx 7 between extensional and compressional stresses in the Pannonian– questionable whether such a distinction is justified as the study of Carpathian basin system of Central Eastern Europe is presented. This Phanerozoic rifts revealed that rift-related volcanic activity and area is the site of pronounced lithospheric strength contrasts between doming of rift zones is basically a consequence of lithospheric exten- the Pannonian Basin area, which is probably underlain by the hottest sion and is not the main driving force of rifting. The fact that rifts can and the weakest lithosphere of continental Europe, the flanking become tectonically inactive at all stages of their evolution, even if Carpathian arc and the particularly strong East-European Platform lith- they have progressed to the Red Sea stage of limited sea-floor osphere. Noteworthy features of the Pannonian Basin system are the spreading (e.g., Bay of Biscay-Pyrenean rift), supports this concept. short duration of the back-arc rifting phases that controlled the collapse However, as extrusion of large volumes of rift-related subalkaline of the internal Carpathian–Dinarides system and the subsequent com- tholeiites must be related to a thermal anomaly within the upper pressional reactivation of this extensional system. This offers a unique mantle, a distinction between ‘active’ and ‘passive’ rifting is to a cer- opportunity to study basin evolution in the aftermath of continental tain degree still valid, though not as ‘black and white’ as originally collision. envisaged. Rifting activity, preceding the breakup of continents is probably 4. Extensional tectonics: concepts and global-scale observations governed by forces controlling the movement and interaction of lithospheric plates. These forces include plate boundary stresses, such as The development of extensional basins and their evolution into slab pull, slab roll-back, ridge push and collisional resistance, and fric- intracontinental thermal sag basins, passive continental margins or tional forces exerted by the convecting mantle on the base of the lith- back-arc basins is controlled by a wide range of parameters, such as osphere (Bott, 1993; Forsyth and Uyeda, 1975; Ziegler, 1993; see also their plate tectonic setting and position relative to collisional plate mar- Wessel and Müller, 2007). gins, the dynamics of the sublithospheric mantle, the thermo-tectonic On the other hand, deviatoric tensional stresses, inherent to the age of the lithosphere, the presence of structural inheritances and the thickened lithosphere of young orogenic belts, as well as those devel- origin and persistence of the controlling stress regime. oping in the lithosphere above upwelling mantle convection cells and mantle plumes (Bott, 1993) do not appear to cause, on their own, the 4.1. Extensional basin systems breakup of continents. However, if such stresses interfere constructively with plate boundary and/or mantle drag stresses, the yield Tectonically active rifts, palaeo-rifts, passive margins and back-arc strength of the lithosphere may be exceeded, thus inducing rifting. basins form a group of genetically related extensional basins that play It must be understood that mantle drag forces are exerted on the an important role in the spectrum of sedimentary basin types (Bally base of a lithospheric plate if its velocity and direction of movement and Snelson, 1980; Ziegler, 1992, 1996b; Ziegler and Cloetingh, 2004; differs from the velocity and direction of the mantle flow. Mantle see also Buck, 2007). Extensional basins cover large areas of the globe drag forces are, however, hard to quantify. Mantle drag stresses at and contain important mineral deposits and energy resources. A the base of the lithosphere cannot be very high (0.05 MPa assuming large number of major hydrocarbon provinces are associated with an asthenosphere viscosity of 5 × 1019 Pa∗s and strain rate on the intracontinental rifts (e.g., North Sea, Sirt and West Siberian basins, order of 10−15 s−1). Consequenly, important drag forces (N1011 N) Dniepr–Donets and Gulf of Suez grabens) and passive margins may be created only in case of very large-scale mantle–lithosphere (e.g., Campos Basin, Gabon, Angola, Mid-Norway and NW Australian interactions, at high deformation rates and when basal stresses are shelves, Niger and Mississippi deltas; Ziegler, 1996a, 1996b). On applied on large areas (hundred thousands of km2)andoverlarge these basins, the petroleum industry has acquired large databases distances (several thousands of km). If these conditions are satisfied, that document their structural styles and allow for a detailed recon- mantle drag can constructively or destructively interfere with plate struction of their evolution. As discussed above, academic geophysi- boundary forces, and thus can either contribute towards plate cal research programs have provided information on the crustal and motion or resist it. Correspondingly, mantle drag can give rise to the lithospheric configurations of tectonically active rifts, palaeo-rifts, build-up of extensional as well as compressional intraplate stresses passive margins and back-arc basins. Petrologic and geochemical studies that contribute to the far-field forces driving lithospheric deformation have advanced the understanding of rift-related magmatic processes. Nu- (Artemieva and Mooney, 2002; Bott, 1993; Forsyth and Uyeda, 1975; merical models, based on geophysical and geological data, have contribut- Warners-Ruckstuhl et al., 2012). Although the present lithospheric ed at lithospheric and crustal scales toward the understanding of dynamic stress field can be readily explained in many areas in terms of plate processes that govern the evolution of extensional basins (Ziegler, 1996b; boundary (far-field) forces (Cloetingh and Wortel, 1986; Richardson, Ziegler and Cloetingh, 2004). 1992; Zoback, 1992), mantle drag probably contributed significantly A natural distinction can be made between tectonically active to the Triassic–Early Cretaceous breakup of Pangea, during which Africa and inactive rifts, and between rifts that evolved in continental and oce- remained nearly stationary and straddled an evolving upwelling and anic lithospheres. Tectonically active intracontinental rifts, such as the radial outflowing mantle convection cell (e.g., Ziegler et al., 2001). How- Rhine Graben, the East African Rift, the Baikal Rift and the Shanxi Rift ever, it appears that with exception of mega-scale events such as forma- of China, are commonly associated with seismicity and volcanic activity tion of spreading centres, the major impact of mantle–lithosphere that pose significant natural hazards. The mid-ocean ridges form an im- interactions refers to normal buoyancy-driven forces and thermo- mense intra-oceanic active rift system that encroaches onto continents mechanical erosion at the base of the lithosphere rather than to basal in the Red Sea and the Gulf of California. Rifts that are tectonically no shear arriving from the mantle drag. Normal forces create upward or longer active are referred to as palaeo-rifts, aulacogens, inactive or downward bending of the lithosphere that may generate important aborted rifts and failed arms, in the sense that they did not progress to concentrations of flexural extensional and compressional stresses crustal separation. Conversely, the evolution of successful rifts culmi- inside the bending plates (e.g., Burov and Diament, 1995). These stresses, nated in the breakup of continents, the opening of new oceanic basins in conjunction with far-field forces, are major factors effecting rift evo- and the development of conjugate pairs of passive margins. lution. Mechanical stretching of the lithosphere and thermal attenua- In the past, a genetic distinction was made between ‘active’ and tion of the lithospheric mantle are associated with the development of ‘passive’ rifting (Olsen and Morgan, 1995; Sengor and Burke, 1978). local deviatoric tensional stresses, which play an increasingly important ‘Active’ rifts are thought to evolve in response to thermal upwelling role during advanced rifting stages (Ziegler, 1993). This has led to the of the asthenosphere (Bott and Kusznir, 1979; Dewey and Burke, development of the concept that many rifts go through an evolutionary 1975), whereas ‘passive’ rifts develop in response to lithospheric ex- cycle starting with an initial ‘passive’ phase that is followed by a more tension driven by far-field stresses (McKenzie, 1978). It is, however, ‘active’ stage during which magmatic processes play an increasingly

Please cite this article as: Cloetingh, S., et al., The Moho in extensional tectonic settings: Insights from thermo-mechanical models, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.06.010 8 S. Cloetingh et al. / Tectonophysics xxx (2013) xxx–xxx important role (Burov and Cloetingh, 1997; Huismans and Beaumont, slab from the lithosphere (Andeweg and Cloetingh, 1998; Bott, 1993); 2011; Huismans et al., 2001a; Wilson and Guiraud, 1992). and (3) retrograde metamorphism of the crustal roots, involving, in Atlantic-type rift systems evolve during the breakup of major conti- the presence of fluids, the transformation of eclogite to less dense nental masses, presumably in conjunction with a reorganization of granulite (Le Pichon et al., 1997; Straume and Austrheim, 1999). the mantle convection system (Ziegler, 1993). During early phases of Although tensional collapse of an orogen can commence shortly after rifting, large areas around future zones of crustal separation can be its consolidation (e.g. the Variscan Orogen, Ziegler, 1990; Ziegler et al., affected by tensional stresses, giving rise to the development of com- 2004), it may be delayed by as much as 30 My, as in the case of the plex and wide graben systems. In time, rifting activity concentrates in Appalachian–Mauretanides (Ziegler, 1990). the zone of future crustal separation, with tectonic activity decreasing Dynamic thermo-mechanical models are important to analyse the and ultimately ceasing in lateral graben systems. In time and as a conse- effect of boundary conditions such as extension velocity and surface quence of progressive lithospheric attenuation and ensuing crustal erosion on the mode of rifting. Extension velocity is a key parameter, doming, local deviatoric tensional stresses play an important secondary governing the focusing of deformation and breakup. In addition, there role in the evolution of such rift systems. Upon crustal separation, are important implications of the lithosphere rheology with respect to the diverging continental margins (pericontinental rifts) and the ‘un- extension — the relationship of viscous flow power-law exponent and successful’ intracontinental branches of the respective rift system be- development of necking (Fletcher, 1974; Schmalholtz, 2008)and come tectonically inactive. However, during subsequent tectonic Moho geometry (e.g., Tirel et al., 2008). The Influence of melting- cycles, such aborted rifts can be tensionally as well as compressionally related weakening and other strain localization processes can also be reactivated (Ziegler et al., 1995, 1998, 2002). Development of Atlantic- considerable. For instance, Huismans and Beaumont (2007) have type rifts is subject to great variations mainly in terms of the duration shown that localization of deformation and rift mode selection during of their rifting stage and the level of volcanic activity (Ziegler, 1988, extension may be closely related to dynamic weakening; some recent 1990). publications emphasize several physical mechanisms to explain this New deep seismic refraction data have been acquired in the recent weakening such as shear heating (Crameri and Kaus, 2010; Kaus and years and these data provide constraints for timing of formation and Podladchikov, 2006), damage evolution (Karrech et al., 2011), geometry of rift basins. It is important to consider in detail the effect grain size reduction (Braun et al., 1999), lattice preferred orientation of magmatism on the Moho geometry, subsidence and the evolution (Tomassi et al., 2009) and some other mechanisms such as fluid- of the lithospheric strength; especially for volcanic rifted margins induced methamorphic reactions at different crustal levels (Mohn such as the Vøring volcanic passive margin (e.g. data papers by et al., 2011). White et al., 2008; Mjelde et al., 2007a, 2007b, 2008; modelling papers by Buck, 2006; Schmeling, 2010; Schmeling and Wallner, 4.1.1. Thermal thinning and stretching of the lithosphere: concepts and 2012; Simon and Podladchikov, 2008). Although continental breakup models might be largely a consequence of plume–lithosphere interaction, Mechanical stretching of the lithosphere can be associated with melting may also develop without invoking a large mantle plume significant magmatic activity and increase in conductive and advective (e.g. van Wijk et al., 2001). In the Northern Atlantic, multi-scale wave- heat flux and consequently an upward displacement of the thermal form inversion of seismic data has resolved many details of both crustal asthenosphere–lithosphere boundary. Small-scale convection in an and upper mantle structures leading to strong evidence for the Iceland evolving asthenospheric diapir may contribute to mechanical thinning Jan Mayen plume system and suggesting a strong impact of mantle of the lithosphere by facilitating lateral ductile mass transfer (Richter dynamics in the North Atlantic region (Rickers et al., 2013). This and McKenzie, 1978). Progressive thermal and mechanical thinning of research demonstrates a close coincidence between the individual the higher-density lithospheric mantle and its replacement by lower- plumes and areas of recent pronounced elevation such as the Southern density asthenosphere induces progressive doming of rift zones. In Scandes of Norway. case of slow rifting, these processes may lead to delamination of the Back-arc rifts are thought to evolve in response to a decrease in lithosphere mantle and hence to accelerated widening of the rift convergence rates and/or even a temporary divergence of colliding (Burov, 2007). At the same time, deviatoric tensional stresses develop- plates, ensuing steepening of the subduction slab and development ing in the lithosphere contribute to its further extension (Bott, 1992). of a secondary upwelling system in the upper plate mantle wedge Although rift-related major hot-spot activity is thought to reflect a above the subducted lower plate lithospheric slab (Uyeda and thermal perturbation of the asthenosphere caused by a deep mantle McCabe, 1983). Changes in convergence rates between colliding plume, smaller-scale ‘plume’ activity may also be the consequence of plates are probably an expression of changes in plate interaction. lithospheric stretching, which is triggered by adiabatic decompression Back-arc rifting can progress to crustal separation and the opening partial melting in areas characterized by an anomalously volatile- of limited oceanic basins (e.g., Sea of Japan, South China Sea, Black rich asthenosphere/lithosphere (White and McKenzie, 1989; Wilson, Sea). However, as convergence rates of colliding plates are variable 1993a). in time, back-arc extensional basins are generally short-lived. Depending on the applicability of the ‘pure-shear’ (McKenzie, 1978) Upon a renewed increase in convergence rates, back-arc extensional or the ‘simple-shear’ model (Wernicke, 1985), or a combination thereof systems are prone to destruction by back-arc compressional stresses (Kusznir et al., 1991), the zone of upper crustal extension, correspond- (e.g., Variscan Rheno–Hercynian Basin, Baikal, Sunda Arc, East China, ing to the subsiding rift, may coincide with the zone of lithospheric Pannonian and Black Sea rift systems; Cloetingh and Kooi, 1992a, mantle attenuation (pure- and combined-shear) or may be laterally off- 1992b; Letouzey et al., 1990; Nikishin et al., 2001; Okay et al., 1994; set from it (i.e. simple-shear). A modification to the pure-shear model is Royden and Horváth, 1988; Thybo and Nielsen, 2009; Uyeda and the ‘continuous depth-dependent’ stretching model which assumes McCabe, 1983; Ziegler, 1990). that stretching of the lithospheric mantle affects a broader area than Extensional disruption of young orogenic belts, involving the the zone of crustal extension (Rowley and Sahagian, 1986). In both development of grabens and pull-apart structures, can be related to models it is assumed that the asthenosphere wells up passively into their postorogenic uplift and the development of deviatoric tensional the space created by mechanical attenuation of the lithospheric mantle. stresses inherent to orogenically overthickened crust (Dewey, 1988; In depth-dependent stretching models this commonly gives rise to Sanders et al., 1999; Stockmal et al., 1986). The following mechanisms flexural uplift of the rift shoulders, and in the flexural cantilever contribute to postorogenic uplift: (1) locking of the subduction zone model, which assumes ductile deformation of the lower crust, this due to decay of the regional compressional stress field (Whittaker produces footwall uplift of the rift flanks and intrabasin fault blocks et al., 1992); (2) roll-back and ultimately detachment of the subducted (Kusznir and Ziegler, 1992). By the same mechanism, the simple-

Please cite this article as: Cloetingh, S., et al., The Moho in extensional tectonic settings: Insights from thermo-mechanical models, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.06.010 S. Cloetingh et al. / Tectonophysics xxx (2013) xxx–xxx 9 shear model predicts asymmetric doming of a rift zone or even flexural isostatic response of the lithosphere to its mechanical stretching and uplift of an arch located to one side of the zone of upper crustal exten- the related thermal perturbation (Van der Beek et al., 1994). sion (Wernicke, 1985). A modification to the simple-shear model The duration of the rifting stage of intracontinental rifts (aborted) envisages that massive upper crustal extensional unloading of the lith- and passive margins (successful rifts) is highly variable (Figs. 5 and 6) osphere causes its isostatic uplift and passive flow of the asthenosphere (Ziegler, 1990; Ziegler and Cloetingh, 2004; Ziegler et al., 2001). Overall, towards the rift axis (Wernicke, 1990). it is observed that, in time, rifting activity concentrates on the zone of The structural style of rifts, as defined at upper crustal and syn-rift future crustal separation with lateral rift systems becoming inactive. sedimentary levels, is influenced by the thickness and thermal state of However, as not all rift systems progress to crustal separation, the dura- the crust and lithospheric mantle at the onset of rifting, by the tion of their rifting stage is obviously a function of the persistence of the amount of crustal extension and the width over which it is distributed, controlling stress field. On the other hand, the time required to achieve the mode of crustal extension (orthogonal or oblique, simple- or crustal separation is a function of the strength (bulk rheology) of the pure-shear) and the lithological composition of the pre- and syn-rift lithosphere, the buildup rate, magnitude and persistence of the exten- sediments (Cloetingh et al., 1995; Ziegler, 1996b). sional stress field, constraints on lateral movements of the diverging A major factor controlling the structural style of a rift zone is the blocks (on-trend coherence, counteracting far-field compressional magnitude of the crustal extensional strain that was achieved across stresses), and apparently less dependent on the availability of pre- it and the distance over which it is distributed (β factor). Although existing crustal discontinuities that can be tensionally reactivated. quantification of the extensional strain and of the stretching factor Crustal separation was achieved in the Liguro–Provençal Basin after is of basic importance for the understanding of rifting processes, 9 My of crustal extension and in the Gulf of California after about 14 My there is often a considerable discrepancy between estimates derived of rifting, whereas opening of the Norwegian–Greenland Sea was from upper crustal extension by faulting, the volume of the rift preceded by an intermittent rifting history spanning some 280 My zone, the crustal configuration and quantitative subsidence analyses (Ziegler, 1988; Ziegler and Cloetingh, 2004). There appears to be (Hinsken et al., 2011; Ziegler and Cloetingh, 2004). Some of these no obvious correlation between the duration of the rifting stage discrepancies can be explained by interplays between surface processes (R) of successful rifts (Fig. 6), which are superimposed on orogenic and subsurface deformation (Burov and Poliakov, 2001)andother belts (Liguro–Provencal Basin, Pyrenees R = 9 My; Gulf of California, factors such as gravitational instabilities in the extending lithosphere Cordillera R = 14 My; Canada Basin, Inuitian fold belt R = 35 My; mantle (Burov, 2007). Central Atlantic, Appalachians R = 42 My; Norwegian–Greenland It is also noteworthy that the above hypotheses and concepts are Sea, Caledonides R = 280 My) and those which developed within based on the inherent assumption of 2D cylindrical (zero strain in the stabilized cratonic lithosphere (southern South Atlantic R = 13 My; out-of-plane direction) deformation. It can be understood that in some northern South Atlantic R = 29 My; Red Sea R = 29 My; BaffinBay cases, the out-of-plane outflow or inflow of mantle–asthenosphere R = 70 My; Labrador Sea R = 80 My). This suggests that the availability material or out-of-plane stress field may significantly affect rift of crustal discontinuities, which regardless of their age (young orogenic evolution. belts, old Precambrian shields) can be tensionally reactivated, does not play a major role in the time required to achieve crustal separation. How- ever, by weakening the crust, such pre-existing or inherited discontinu- 4.1.2. Syn-rift subsidence and duration of rifting stage ities play a role in the localization and distribution of crustal strain. The balance of two mechanisms controls the syn-rift subsidence of Moreover, by weakening the lithosphere, they contribute to the preferen- a sedimentary basin. First, elastic/isostatic adjustment of the crust to tial tensional reactivation of young as well as old orogenic belts (Janssen stretching of the lithosphere and its adjustment to sediment loading et al., 1995; Ziegler et al., 2001). causes subsidence of the mechanically thinned crust (Keen and Boutilier, 1990; McKenzie, 1978). Depending on the depth of the lithospheric necking level, this is accompanied by either flexural uplift or 4.1.3. Post-rift subsidence down warping of the rift zone (Fig. 4)(Braun and Beaumont, 1989; Kooi Similar to the subsidence of oceanic lithosphere, the post-rift subsi- et al., 1992). Second, uplift of a rift zone is caused by upwelling of the dence of extensional basins is mainly governed by thermal relaxation asthenosphere into the space created by mechanical stretching of and contraction of the lithosphere, resulting in a gradual increase of the lithosphere, thermal upward displacement of the asthenosphere– its density and its flexural strength, and by its isostatic response to lithosphere boundary, thermal expansion of the lithosphere and sedimentary loading. Theoretical considerations indicate that subsi- intrusion of melts at the base of the crust (Turcotte and Emerman, dence of post-rift basins follows an asymptotic curve, reflecting the 1983). Thus, the geometry of a rifted basin is a function of the elastic/ progressive decay of the rift-induced thermal anomaly, the magnitude

Fig. 4. Concept of lithospheric necking. The level of necking is defined as the level of no vertical motions in the absence of isostatic forces. Left panel: kinematically induced configuration after rifting for different necking depths. Right panel: subsequent flexural isostatic rebound.

Fig. 5. Duration of rifting stage of ‘abortive’ rifts (paleo-rifts, failed arms). Horizontal bars in My; numbers below horizontal bars indicate onset and termination of rifting stage in My; numbers besides R left of each bar give the duration of rifting stage in My; stars indicate periods of main volcanic activity; D indicates periods of doming. From Ziegler and Cloetingh (2004). of which is thought to be directly related to the lithospheric stretching The shape and dimension of rift-induced asthenosphere–lithosphere value (Beaumont et al., 1982; McKenzie, 1978; Royden et al., 1980; boundary anomalies essentially control the geometry of the evolving Steckler and Watts, 1978; Watts et al., 1982). During the post-rift post-rift thermal-sag basin. Thermal sag basins associated with aborted evolution of a basin, the thermally destabilized continental lithosphere rifts are broadly saucer-shaped and generally overstep the rift zone, re-equilibrates with the asthenosphere (McKenzie, 1978; Steckler and with their axes coinciding with the zone of maximum lithospheric Watts, 1978; Wilson, 1993b). In this process, in which the temperature attenuation. Pure-shear-dominated rifting gives rise to the classical regime of the asthenosphere plays an important role (ambient, below, ‘steer's head’ configuration of the syn- and post-rift basins (White and or above ambient), new lithospheric mantle, consisting of solidified McKenzie, 1989) in which both basin axes roughly coincide, with the asthenospheric material, is accreted to the attenuated old continental post-rift basin broadly overstepping the rift flanks. This geometric lithospheric mantle (Ziegler et al., 1998). On the other hand, lateral relationship between syn- and post-rift basins is frequently observed heat exchanges and progressive increase of the flexural strength of (e.g., North Sea Rift, Ziegler, 1990; West Siberian Basin, Artyushkov the lithosphere with cooling may significantly retard it subsidence and Bear, 1990;Dniepr–Donets Graben, Kusznir et al., 1996,Gulfof (compared to the McKenzie model) after the first 10–20 Ma of post- Thailand, Hellinger and Sclater, 1983; Sudan rifts, McHargue et al., rift subsidence (Burov and Poliakov, 2001). In addition, densification 1992). Such a geometric relationship is compatible with discontinuous, of the continental lithosphere involves crystallization of melts that ac- depth-dependent stretching models that assume that the zone of crust- cumulated at its base or were injected into it, subsequent thermal con- al extension is narrower than the zone of lithospheric mantle attenua- traction of the solidified rocks and, under certain conditions, their phase tion. The degree to which a post-rift basin oversteps the margins of transformation to eclogite facies. The resulting negative buoyancy effect the syn-rift basin is a function of the difference in the width of the is the primary cause of post-rift subsidence. However, in a number of zone of crustal extension and the zone of lithospheric mantle attenua- basins, significant departures from the theoretical thermal subsidence tion (White and McKenzie, 1988) and the effective elastic thickness curve are observed. These can be explained as effects of compressional (EET) of the lithosphere (Watts et al., 1982). intraplate stresses and related phase transformations (Cloetingh and Conversely, simple-shear dominated rifting gives rise to a lateral off- Kooi, 1992b; Lobkovsky et al., 1996; Van Wees and Cloetingh, 1996). set between the syn- and post-rift basin axes. An example is the Tucano

Fig. 6. Duration of rifting stage of ‘successful’ rifts. Legend same as Fig. 5. Numbers beside letter ‘S’ indicate duration of seaﬂoor spreading stage in million years. From Ziegler and Cloetingh (2004).

Graben of north-eastern Brazil, which ceased to subside at the end of induced by rifting that did not progress to crustal separation depends the rifting stage, whereas the coastal Jacuipe–Sergipe– Alagoas Basin, on the magnitude of crustal stretching (δ-factor) and lithospheric man- to which the former was structurally linked, was the site of crustal tle attenuation (β-factor), the thermal regime of the asthenosphere, the and mantle-lithospheric thinning culminating in crustal separation volume of melts generated and whether these intruded the lithosphere and subsequent major post-rift subsidence (Chang et al., 1992; Karner and influenced the evolution of the Moho. After 60 My, about 65%, and et al., 1992). Moreover, the simple-shear model can explain the after 180 My about 95% of a deep-seated thermal anomaly associated frequently observed asymmetry of conjugate passive margins and with a major pull-up of the asthenosphere–lithosphere boundary differences in their post-rift subsidence pattern (e.g., Central Atlantic have decayed (mantle-plume model). and Red Sea: Favre and Stampfli, 1992). Discrepancies in the post-rift The thickness of the post-rift sedimentary column that can accumu- subsidence of conjugate margins are attributed to differences in their late in passive margin basins and in thermal-sag basins above aborted lithospheric configuration at the crustal separation stage. At the end rifts is not only a function of the magnitude of the lithospheric mantle of the rifting stage, a relatively thick lithospheric mantle supports and crustal attenuation factors, but also of the population of upper lower plate margins, whereas upper plate margins are underlain by crustal faults (Ziegler, 1983, 1990). This discrepancy is somewhat a strongly attenuated lithospheric mantle and partly directly by the as- reduced when flexural isostasy is assumed (Roberts et al., 1993). thenosphere. Correspondingly, lower plate margins are associated at However, because during rifting attenuation of the lithosphere is not the crustal separation stage with smaller thermal anomalies than only achieved by its mechanical stretching but also by convective upper plate margins (Stampfli et al., 2001; Ziegler et al., 1998). These and thermal upward displacement of the asthenosphere–lithosphere differences in lithospheric configuration of conjugate simple-shear boundary, the magnitude of a thermal anomaly derived from the margins have repercussions on their rheological structure, even after post-rift subsidence of a basin cannot be directly related to a mechanical full thermal relaxation of the lithosphere, and their compressional stretching factor. Moreover, intraplate compressional stresses and reactivation potential (Ziegler et al., 1998). phase transformations in the lower crust and lithospheric mantle have The magnitude of post-rift tectonic subsidence of aborted rifts and an overprinting effect on post-rift subsidence and can cause significant passive margins is a function of the thermal anomaly that was intro- departures from a purely thermal subsidence curve (Cloetingh and duced during their rifting stage and the degree to which the lithospheric Kooi, 1992b; Lobkovsky et al., 1996; Van Wees and Cloetingh, 1996). mantle was thinned. Most intense anomalies develop during crustal Furthermore, during rifting the ascent of mantle-derived melts to separation, particularly when a plume assisted and asthenospheric the base of the crust can cause destabilization of the Moho, magmatic melts well up close to the surface. The magnitude of thermal anomalies inflation of the crust and its metasomatic reactivation and secondary

Please cite this article as: Cloetingh, S., et al., The Moho in extensional tectonic settings: Insights from thermo-mechanical models, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.06.010 12 S. Cloetingh et al. / Tectonophysics xxx (2013) xxx–xxx differentiation (Mohr, 1992; Morgan and Ramberg, 1987; Stel et al., the lithosphere inherited from the syn-rift phase. Moreover, horizontal 1993; Watts and Fairhead, 1997). For instance, lower crustal velocities stresses in the lithosphere strongly affect the development of salt of 6.30–7.2 km/s and densities of 3.02 × 103 kg/m3 characterize the diapirism, accounting for local subsidence anomalies (Cloetingh and highly attenuated crust of the Devonian Dniepr–Donets rift (Fig. 1B) Kooi, 1992b), and have a strong impact on the hydrodynamic regime almost up to the base of its syn-rift sediments, probably owing to its of extensional basins (Van Balen and Cloetingh, 1994) by contributing syn-rift permeation by mantle-derived melts (Maystrenko et al., 2003; to the development of overpressure, as seen in parts of the Pannonian Yegorova et al., 1999). Basin (Van Balen et al., 1995). Glacial loading and unloading can further Moreover, accumulation of very thick post-rift sedimentary complicate post-rift lithospheric motions, the scope of which requires sequences can cause phase transformation of crustal rocks to granulite further analysis (Solheim et al., 1996) (e.g. post-glacial rebound of the facies, with the resulting densification of the crust accounting for accel- Barents Sea shelf). erated basin subsidence. This process can be amplified in the cases of These considerations indicate that stretching factors derived from eclogite transformation of basaltic melts that were injected into the the subsidence of post-rift basins must be treated with reservations. lithospheric mantle or that had accumulated at its base (Lobkovsky Nevertheless, quantitative subsidence analyses, combined with other et al., 1996). However, it is uncertain whether large-scale eclogite for- data, are essential for the understanding of post-rift subsidence pro- mation can indeed occur at crustal thicknesses of 30–40 km (Carswell, cesses by giving a measure of the lithospheric anomaly that was intro- 1990; Griffinetal.,1990). Although phase transformations, entailing duced during the rifting stage of a basin and by identifying deviations an upward displacement of the geophysical Moho, are thought to from purely thermal cooling trends. occur under certain conditions at the base of very thick Proterozoic cratons (increased confining pressure due to horizontal intraplate 4.1.4. Post-rift compressional reactivation of extensional basins stresses and/or ice load, possible cooling of asthenosphere, Cloetingh Palaeo-stress analyses give evidence for changes in the magnitude and Kooi, 1992b), it is uncertain whether physical conditions conducive and orientation of intraplate stress fields on time scales of a few mil- to such transformations can develop in response to post-rift sedimenta- lion years (Bergerat, 1987; Dèzes et al., 2004). Thus, in an attempt to ry loading of palaeo-rifts and passive margins as, for example, suspected understand the evolution of a post-rift basin, the effects of tectonic for the East Newfoundland Basin (Cloetingh and Kooi, 1992b) and the stresses on subsidence must be separated from those related to ther- West Siberian Basin (Lobkovsky et al., 1996). mal relaxation of the lithosphere (Cloetingh and Kooi, 1992b). An alternative process to account for the development of post-rift In response to the buildup of far-field compressional stresses, rifted sag basins in passive margins is the lateral viscous flow operating in basins, characterized by a strongly faulted and thus permanently the lower crust away from the main depocenters. Lateral viscous weakened crust, are prone to reactivation at all stages of their post-rift flow in the lower crust is also well evidenced in the core complexes evolution, resulting in their inversion (Ziegler, 1990; Ziegler et al., associated with the collapse of orogens (i.e. in the North American 1995, 1998, 2001). Only under special condition can gravitational forces Cordillera), and at early stages of back-arc opening, prior to the for- associated with topography around a basin cause its inversion (Bada mation of true oceanic lithosphere (i.e. in the Aegean back arc and et al., 2001; Pascal and Cloetingh, 2009). Algerian Basin, the later being still characterized by a layered lower Rheological considerations indicate that the lithosphere of ther- crust that could represent the remnants of hyper-extended continen- mally stabilized rifts, lacking a thick post-rift sedimentary prism, is tal material (Roure et al., 2009). considerably stronger than the lithosphere of adjacent unstretched As rifting processes can take place intermittently over very long areas (Ziegler and Cloetingh, 2004). This contradicts the observation periods of time (e.g., Norwegian-Greenland Sea rift, Ziegler, 1988), that rift zones and passive margins are preferentially deformed during thermal anomalies introduced during early stretching phases start periods of intraplate compression (Ziegler et al., 2001). However, burial to decay during subsequent periods of decreased extension rates. of rifted basins under a thick post-rift sequence contributes by thermal Thus, the thermal anomaly associated with a rift may not be at its blanketing to weakening of their lithosphere (Stephenson, 1989), thus maximum when crustal stretching terminates and the rift becomes rendering it prone to tectonic reactivation. inactive. Similarly, late rifting pulses and/or regional magmatic events may interrupt and even reverse lithospheric cooling processes 4.1.5. Finite strength of the lithosphere during extensional basin (e.g., Palaeocene thermal uplift of northern parts of Viking Graben due formation to Iceland plume impingement: Nadin and Kusznir, 1995; Ziegler, In recent approaches to extensional basin modelling the implemen- 1990). Therefore, analyses of the post-rift subsidence of rifts that have tation of finite lithospheric strengths has been an important step evolved in response to multiple rifting phases spread over a long period forward. Advances in the understanding of lithospheric mechanics need to take their entire rifting history into account. (Burov, 2011) demonstrate that early stretching models, which Intraplate compressional stress, causing deflection of the lithosphere assumed zero lithospheric strength during rifting, are not valid in response to lithospheric folding, can seriously overprint the thermal (e.g., Bassi, 1995; Chery et al., 1992). Moreover, the notion of a possible subsidence of post-rift basins (see for a recent discussion Cloetingh decoupling zone between the strong upper crust and the strong litho- et al., 2008). The example of the Plio-Pleistocene evolution of the spheric upper mantle is also important in the context of extensional North Sea shows that the buildup of regional compressional stresses basin formation. During the past decades the relative importance of can cause a sharp acceleration of post-rift subsidence (Cloetingh et al., the pure shear (McKenzie, 1978) and simple shear (Wernicke, 1985) 1990; Kooi et al., 1991; Van Wees and Cloetingh, 1996). Similar contem- mode of extension has been a matter of debate. In the presence of a poraneous effects are recognized in North Atlantic passive-margin basins weak lower crustal layer, decoupling of the mechanically strong upper (Cloetingh et al., 1990) as well as in the Pannonian Basin (Horváth and mantle from the strong upper crust depends on their thermal age and Cloetingh, 1996).AlsothelateEoceneacceleratedsubsidenceofthe the thickness of their sedimentary cover. Black Sea Basin can be attributed to the buildup of a regional compres- During the last few years, the role played by rheology during sional stress field (Cloetingh et al., 2003; Robinson et al., 1995). extensional basin formation was gradually appreciated due to the At present, many post-rift extensional basins are in a state of hori- availability of improved modelling, increasingly shifting attention zontal compression as documented by stress indicators summarized away from these end lithospheric members. Finite element models in the World Stress Map (Heidbach et al., 2010). The magnitude of permitted to explore the large-scale implications of a finite lithospheric stress-induced vertical motions of the lithosphere during the post-rift strength and particularly its sensitivity to the presence of fluids in the phase, causing accelerated basin subsidence and tilting of basin crust and lithospheric mantle (Bassi, 1995; Braun and Beaumont, margins, depends on the ratio of the stress level and the strength of 1989; Dunbar and Sawyer, 1989). These differences are a direct

Please cite this article as: Cloetingh, S., et al., The Moho in extensional tectonic settings: Insights from thermo-mechanical models, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.06.010 S. Cloetingh et al. / Tectonophysics xxx (2013) xxx–xxx 13 consequence of contrasts in thermo-tectonic age between young conti- the distribution of fission-track length permits to backstack the eroded nental versus and old cratonic lithosphere affected by extensional basin sediments from their present position in the basin to their source on formation. These dynamic models, which require intensive and expen- the rift shoulder in an effort to obtain a better reconstruction of the sive computing, are not always suitable for a user-oriented industry en- rift-shoulder geometries (e.g., Burov and Cloetingh, 1997; Redfield vironment. However, they provided the background for a more user- et al., 2005; Rohrman et al., 1995; Van der Beek et al., 1995). This has friendly class of kinematic models targeted at modelling rift-shoulder led to a better understanding of the timing and magnitude of rift- uplift and the basin fill. For the development of extensional basins shoulder uplift. these kinematic models invoke necking of the lithosphere around one A second research line focused on the development of a model for ofitsstronglayers(seeBraun and Beaumont, 1989; Kooi et al., 1992; basin fill simulation, integrating the effects of rift-shoulder erosion Spadini et al., 1995). Fig. 4 illustrates the basic features of these models through hill-slope transport and river incision with sediment deposition and their relation to the strength distribution within the lithosphere. In in the basin. These models (e.g., van Balen et al., 1995), predict the the presence of a strong layer in the subcrustal mantle, the lithospheric progradation of sedimentary wedges into extensional basins and the necking level is deep, inducing pronounced rift-shoulder uplift. This development of hinterland basins having a distinctly different strati- type of response is to be expected if extension affects cold and corre- graphic signature than predicted by standard models, which invoke spondingly strong intracontinental lithosphere with an average crustal stretching and post-rift flexure, as commonly applied in the software thickness of 30–40 km; it is commonly observed in intracratonic rifts packages. Testing this model against a number of rifted margins around and rifted margin such as, for example, the Red Sea and the Trans- the world demonstrates that erosion of the rift-shoulder topography, Antarctic Mountains (Chery et al., 1992; Cloetingh et al., 1995). created during extension of a lithosphere with a finite strength, can For Alpine/Mediterranean basins, which developed on young, sometimes eliminate the need to invoke eustatic sea-level changes to orogenically thickened crust and lithosphere, the necking level is explain large-scale stratigraphic features of rifted basins and associated generally located at depths of 5 to 10 km within the crust (Fig. 4). An hinterland basins. The tectonically-driven erosion of these rift shoulders example of such a situation is found in the Pannonian Basin (Horváth is commonly observed on the flanks of extensional basins such as the and Cloetingh, 1996; Van Balen et al., 1995) where the strength of the Catalan Coastal Ranges facing the Neogene Valencia trough of the lithospheric upper mantle is almost zero. Important exceptions to this Western Mediterranean (Gaspar-Escribano et al., 2003), or the large general pattern do occur, however, such as in the Southern Tyrrhenian scale exhumation of the onshore Morroccan segment of the Atlantic Sea, for which our modelling indicates a deep necking level to fit obser- margin (Gouiza et al., 2010). vational data (Spadini et al., 1995). This is primarily attributed to the Considering the notion that in syn-rift basins different spatial and fact that the Southern Tyrrhenian Basin developed essentially on temporal scales are by their very nature linked, the lack of constraints Hercynian lithosphere with significant bulk strength of its mantle on the evolution and structural configuration of surrounding areas component. severely limits modelling of syn-rift sedimentary sequences. More- Similarly, models analyzing the differences between the western over, modelling can assess tectonic controls on syn-rift depositional and eastern sub-basins of the Black Sea have inferred important sequences at a sub-basin scale (see also Nottvedt et al., 1995). The mag- variations in lithospheric necking depth and thermal conditions related nitude of footwall uplift of individual fault blocks appears to be a func- to the mode and timing of extension, as well as the thermo-tectonic age tion of the lithospheric necking level and thus cannot be attributed to of the underlying lithosphere (Cloetingh et al., 2003; Robinson et al., factors restricted to the sub-basin scale. 1995; Spadini et al., 1996, 1997). The inter-dependence of the necking depth and pre-rift crustal and lithospheric thicknesses, the effective 5. Rheological stratification of the lithosphere and basin evolution elastic thickness (EET) of the lithosphere and strain rates was inves- tigated in a comparative study on several basins, which are all 5.1. Lithosphere strength and deformation mode superimposed on Hercynian or Alpine destabilized crust (Cloetingh et al., 1995). Results of this study suggest that for Alpine/Mediterranean The major, directly measurable proxy for the integrated strength basins the position of the necking level depends primarily on the of the lithosphere is its EET, (or Te, see Watts, 2001 and references pre-rift crustal thickness and strain rate, whereas the key controlling therein). By comparing observations of flexure in the region of long- factors in intra-cratonic rifts, which are superimposed on Precambrian term loads such as ice, sediment and volcanoes to the predictions of crust, appear to be the pre-rift lithospheric thickness and strain rate. elastic plate flexure models, Te of the lithosphere in a wide range of These models illustrate the importance of the pre-rift lithosphere rheol- geological settings could be estimated. Flexure studies on oceanic ogy for the architecture of the evolving extensional basins and the asso- lithosphere suggest that its Te ranges between 2 and 40 km and ciated pattern of vertical motions. Moreover, they demonstrate that the depends on plate age and sedimentary load. The Te of continental litho- better the pre-rift evolution of an area is constrained, the greater the sphere ranges from 0 to 100 km and shows no clear relationship with chance to define the proper parameters for successful modelling and the consolidation age of the crust (Watts, 2001). The results of these the definition of the underlying syn-rift mechanisms. flexure studies are qualitatively consistent with the results of experimental rock mechanics. The Brace–Goetze failure envelope curves 4.1.6. Rift-shoulder development and architecture of basin fill (Brace and Kohlstedt, 1980; Goetze and Evans, 1979), for example, As discussed above, the finite strength of the lithosphere has predict that strength increases with depth and then decreases in accor- important implications for the crustal structure of extensional basins dance with the brittle (e.g. Byerlee, 1978) and ductile deformation laws. and the development of accommodation space within them. The In oceanic regions, the envelopes are approximately symmetric about development of significant rift-shoulder topography in response to the depth of the Brittle–Ductile Transition (BDT) where the brittle– lithospheric extension has drawn attention to the need to constrain elastic and elastic–ductile layers contribute equally to the strength. the coupled vertical motion of the uplifting rift shoulders and the Furthermore, the lithospheric strength is also dependent on rock subsiding basin (Chery et al., 1992; Kusznir and Ziegler, 1992). Whereas composition and fluid content (Burov, 2011). the standard approach in basin analysis focused until recently primarily The strength of continental lithosphere is controlled by its stratified on the subsiding basin, treating sediment supply as an independent pa- depth-dependent rheological structure in which the thickness and rameter, necking models highlight the need for linking sediment supply composition of the crust, the thickness of the lithospheric mantle, the to the rift-flank uplift and their erosion history. To this end, a two-fold potential temperature of the asthenosphere, and the presence or approach was followed. The first research line aimed at constraining absence of fluids, as well as strain rates play a dominant role. By con- the predicted uplift histories by geothermochronology. Modelling of trast, the strength of oceanic lithosphere depends on its thermal regime,

Te values observed due to wide variations in crustal thickness and as the crustal fabric and probably are related to subducted oceanic composition, and geothermal gradients. and/or continental crustal material, suggests that the continental It has been also suggested (Jackson, 2002) that the observed maxi- lithospheric mantle is not necessarily homogenous but can contain mal depth of earthquake occurrence (Ts) correlates with Te and hence lithological discontinuities that enhance its mechanical anisotropy is also a proxy for the integrated strength of the lithosphere. From (Vauchez et al., 1998; Ziegler et al., 1998). Such discontinuities, observations it is evident that mantle seismicity is rare in continents, consisting of eclogitized crustal material, can potentially weaken suggesting that their lithospheric strength is concentrated in the conti- the strong upper part of the lithospheric mantle. Moreover, even in nental crust (so called “crème-brûlé” rheology model, Burov and Watts, the face of similar crustal thicknesses, the heat ﬂow of deeply degraded 2006; Handy and Brun, 2003). These concepts were, however, later Late Proterozoic and Phanerozoic orogenic belts is still elevated as dismissed (Audet and Bürgmann, 2011; Burov and Watts, 2006; Watts compared to adjacent old cratons (e.g., Pan African belts of Africa and and Burov, 2003) on account of (1) the lack of physical grounds for a Arabia; Janssen, 1996). This is probably due to the younger age of

Te − Ts correlation (2) continental Te values being often larger than Ts, their lithospheric mantle and possibly also to a higher radiogenic heat and (3) the rare occurrence of sub-Moho earthquakes in continents. generation potential of their crust. These factors contribute to weakening However, the rarity of sub-Moho seismicity still needs to be explained of former mobile zones to the end that they present rheologically weak even though some first order explanations have been advanced such zones within a craton, as evidenced by their preferential reactivation as an increase of the brittle strength with depth, domination of grain during the breakup of Pangea (Janssen et al., 1995; Ziegler, 1989a, boundary sliding creep flow in case of lithospheric mantle deformation. 1989b; Ziegler et al., 2001). The strength of continental crust depends largely on its composition, From a rheological point of view, the thermally destabilized thermal regime and the presence of fluids, and also on the availability of lithosphere of tectonically active rifts, as well as of rifts and passive pre-existing crustal discontinuities (Ziegler and Cloetingh, 2004;see margins that have undergone only a relatively short post-rift evo- also Burov, 2011). Deep-reaching crustal discontinuities, such as lution (e.g., 25 Ma), is considerably weaker than that of thermally thrust- and wrench-faults, cause significant weakening of the otherwise stabilized rifts and of unstretched lithosphere (Ziegler et al., mechanically strong upper parts of the crust. As such discontinuities are 1998). In this respect, it must be realized that, during rifting, apparently characterized by a reduced frictional angle, particularly in progressive mechanical and thermal thinning of the lithospheric the presence of fluids, they are prone to reactivation at stress levels mantle and its substitution by the upwelling asthenosphere is ac- that are well below those required for the development of new faults. companied by a rise in geotherms causing progressive weakening Deep reflection-seismic profiles show that the crust of Late Proterozoic of the extended lithosphere. In addition, its permeation by fluids and Palaeozoic orogenic belts is generally characterized by a monoclinal causes its further weakening. Upon decay of the rift-induced ther- fabric that extends from upper crustal levels down to the Moho at mal anomaly, rift zones are, rheologically, considerably stronger which it either soles out or by which it is truncated. This fabric reflects than unstretched lithosphere. However, accumulation of thick the presence of deep-reaching lithological inhomogeneities and shear syn- and post-rift sedimentary sequences can cause by thermal zones (Ziegler and Cloetingh, 2004). blanketing a weakening of the strong parts of the upper crust and The strength of the continental upper lithospheric mantle de- lithospheric mantle of rifted basins (Stephenson, 1989). Moreover, pends to a large extent on the thickness of the crust but also on as faults may permanently weaken the crust of rifted basins due to its age and thermal regime (see Jaupart and Mareschal, 2007; strain softening mechanisms activated during their formation Toussaint et al., 2004). Thermally stabilized stretched continental (Huismans and Beaumont, 2003), they are prone to tensional as lithosphere with a 20 km thick crust and a lithospheric mantle well as compressional reactivation (Ziegler et al., 1995, 1998, thickness of 50 km is mechanically stronger than unstretched litho- 2001, 2002). sphere with a 30 km thick crust and a 70 km thick lithospheric mantle. In view of its rheological structure, the continental lithosphere can Extension of stabilized continental crustal segments precludes ductile be regarded under certain conditions as a two-layered viscoelastic flow of the lower crust and faults will be steep to listric and propagate beam (Reston, 1990; ter Voorde et al., 1998). The response of such a sys- towards the hanging wall, that is, towards the basin centre (Bertotti et tem to the buildup of extensional and compressional stresses depends al., 2000). on the thickness, strength and spacing of the two competent layers, Under these conditions, the lower crust will deform by distributing on stress magnitudes and strain rates and the thermal regime (Watts ductile shear in the brittle–ductile transition domain. This is compatible and Burov, 2003; Zeyen et al., 1997). As the structure of continental with the occurrence of earthquakes within the lower crust and even lithosphere is also areally heterogeneous, its weakest parts start to close to the Moho (e.g., southern Rhine Graben: Bonjer, 1997; East yield first, once tensional or compressional intraplate stress levels African rifts: Shudofsky et al., 1987). equate their strength.

Fig. 7. Compilation of observed and predicted values of effective elastic thickness (EET), depth to bottom of mechanically strong crust (MSC), and depth to bottom of mechanically strong lithospheric mantle (MSL) plotted against the age of the continental lithosphere at the time of loading and comparison with predictions from thermal models of the lithosphere. Labelled contours are isotherms. Isotherms marked by ‘solid lines’ are for models that account for additional radiogenic heat production in the upper crust. ‘Dashed lines’ correspond to pure cooling models for continental lithosphere. The equilibrium thermal thickness of the continental lithosphere is 250 km. The yellow band corresponds to depth intervals marking the base of the mechanical crust (MSC) while the blue band marks the strong part of the mantle lithosphere (MSL). ‘Squares’ correspond to EET estimates, ‘circles’ indicate MSL estimates, and ‘diamonds’ correspond to estimates of MSC. ‘Bold letters’ correspond to directly estimated EET values derived from ﬂexural studies on, for example, foreland basins, ‘Thinner letters’ indicate indirect rheological estimates derived from extrapolation of rock-mechanics studies. The data set includes (I): Old thermo-mechanical ages (1000–2500 Ma): northernmost (N.B.S.), central (C.B.S.), and southernmost Baltic Shield (S.B.S.); Fennoscandia (FE); Verkhoyansk plate (VE); Urals (UR); Carpathians; Caucasus; (II): Intermediate thermo-mechanical ages (500–1000 Ma): North Baikal (NB); Tarim and Dzungaria (TA-DZ); Variscan of Europe: URA, NHD, EIFEL; and (III): Younger thermo-mechanical ages (0–500 Ma): Alpine belt: JURA, MOLL (Molasse), AAR; southern Alps (SA) and eastern Alps (EA); Ebro Basin; Betic rifted margin; Betic Cordilleras. Modiﬁed from Cloetingh and Burov (1996).

5.2. Mechanical controls on the evolution of rifts: Europe's continental and Burov, 1996; Cloetingh et al., 2005b; Perez-Gussinye and Watts, lithosphere 2005). An understanding of the temporal and spatial strength distribution in the NW European lithosphere may offer quantitative The thermo-mechanical age concept provides the framework for insights into the patterns of its intraplate deformation (basin inversion, EET estimates of the lithosphere. Fig. 7 was constructed on the basis up-thrusting of basement blocks), and particularly into the pattern of of a large data set for Eurasian foreland basins (Cloetingh and lithosphere-scale folding, as described by Ziegler et al. (1995, 1998). Burov, 1996) and illustrates the general trend of increasing elastic Owing to the large amount of high-quality geophysical data ac- plate thickness with increasing thermo-mechanical age of the conti- quired during the last 20 years in Europe, its lithospheric conﬁguration nental lithosphere. Moreover, Fig. 7 summarizes the bulk rheology is rather well known; though signiﬁcant uncertainties remain in many of the lithosphere based on extrapolations from rock mechanical areas concerning the seismic and thermal thickness of the lithosphere data, constrained by crustal geophysical data and thermal models. A (Artemieva, 2006, 2011a, 2011b; Artemieva and Mooney, 2001; characteristic feature of these models is the incorporation of a Babuska and Plomerova, 1992). Nevertheless, available data help to quartz-dominated upper crustal rheology and an olivine-controlled constrain the rheology of the European lithosphere, thus enhancing mantle rheology. These models were designed to shed light on the our understanding of its strength. depth to the base of the mechanically strong upper part of the crust Strength envelopes and the effective elastic thickness of the litho- (MSC) and the mechanically strong part of the upper mantle sphere have been commonly calculated for a number of locations in lithosphere (MSL). Analysis of Fig. 7 demonstrates that the mechanical Europe (e.g., Cloetingh and Burov, 1996). However, as such calculations properties of the crust are little affected by its age-dependent cooling, were made for scattered points only, or along transects, they provide whereas the thickness and strength of the lithospheric mantle is very limited information on lateral strength variations of the lithosphere. strongly age dependent. Depth- and temperature-dependent rheologi- Although lithospheric thickness and strength maps have already been cal models show that the mechanically weak lower crust separates constructed for the Pannonian Basin (Lankreijer et al., 1999) and the the mechanically strong upper crust and upper lithospheric mantle Baltic Shield (Kaikkonen et al., 2000), such maps were until recently (e.g., Watts and Burov, 2003). The EET bands for the mechanically not yet available for all of Europe. strong upper crust and upper lithospheric mantle describe the integrat- As evaluation and modelling of the response of the lithosphere to ed EET of the continental lithosphere that has a bearing on its response vertical and horizontal loads (e.g., Toussaint et al., 2004)requiresan to loads imposed on it. The degree of coupling and/or decoupling understanding of its strength distribution, dedicated efforts were between these two mechanically strong layers plays an important role made to map the strength of the European lithosphere by implementing in the structural style compressionally deformed extensional basins. 3-D strength calculations (Cloetingh et al., 2005b; Tesauro et al., 2009). The bulk-strength of the heterogeneous European continental Strength calculations of the lithosphere depend primarily on its lithosphere is directly linked to its thermo-tectonic age (Cloetingh thermal and compositional structure and are particularly sensitive to

Fig. 8. From crustal thickness (top left) and thermal structure (top right) to lithospheric strength (bottom): conceptual conﬁguration of the thermal structure and composition of the lithosphere, adopted for the calculation of 3-D strength models. Modiﬁed from Cloetingh et al. (2005b).

thermal uncertainties (Burov, 2011; Burov and Diament, 1995; Ranalli, Proterozoic lithosphere of the East-European Platform to the east of 1995). For this reason, the workflow aimed at the development of a 3-D the Teisseyre–Tornquist line and the relatively weak Phanerozoic litho- strength model for Europe was twofold: (1) construction of a 3-D com- sphere of Western Europe. A similar strength contrast occurs at the positional model and (2) calculating a 3-D thermal cube. The final 3-D transition from strong Atlantic oceanic lithosphere to the relatively strength cube was obtained by calculating 1-D strength envelopes for weak continental lithosphere of Western Europe. Within the Alpine each lattice point (x, y) of a regularized raster covering NW Europe foreland, pronounced NE–SE trending weak zones are recognized that (Fig. 8). For each lattice point, the appropriate input values were coincide with such major geologic structures as the Rhine Graben obtained from a 3-D compositional and thermal cube. A geological System and the North Danish–Polish Trough, which are separated by and geophysical geographic database was used as reference for the the high-strength North German Basin and the Bohemian Massif. More- construction of input models. over, a broad zone of weak lithosphere characterizes the Massif Central For continental realms, a 3-D multilayer compositional model was and surrounding areas. The presence of thickened crust in the area constructed, consisting of one mantle-lithosphere layer, 2–3 crustal of the Teisseyre–Tornquist suture zone (Fig. 9) gives rise to a pro- layers and an overlying sedimentary cover layer, whereas for oceanic nounced mechanical weakening of the lithosphere, particularly of its areas a one-layer model was adopted. For the depth to the different mantle part. interfaces several regional or European-scale compilations were avail- Whereas the lithosphere of Fennoscandia is characterized by a rela- able that are based on deep seismic reflection and refraction or surface tively high strength, the North Sea rift system corresponds to a zone of wave dispersion studies (e.g., Artemieva et al., 2006; Blundell et al., weakened lithosphere. Other areas of high lithospheric strength are the 1992; Suhadolc and Panza, 1989). Information on the depth to the Bohemian Massif and the London–Brabant Massif both of which exhibit base of the crust, was mainly derived from the European Moho map low seismicity. A pronounced contrast in strength can also be noticed by Tesauro et al. (2008) (Fig. 9). Regional compilation maps of the seis- between the strong Adriatic indenter and the weak Pannonian Basin mological lithosphere thickness were used as reference base for the area (see also Fig. 10). thickness of the thermal lithosphere in thermal modelling (Babuska The lateral strength variations of Europe's intraplate lithosphere are and Plomerova, 1993, 2001; Plomerova et al., 2002). primarily caused by variations in the mechanical strength of the litho- Fig. 10 shows spatial variations in integrated strength and effective spheric mantle, whereas variations in crustal strength appear to be elastic thickness of Europe's lithosphere inferred from forward rheolog- more modest. Variations in lithospheric mantle strength are primarily ical modelling (from Tesauro et al., 2009). As is evident from Fig. 10,the related to variations in the thermal structure of the lithosphere that percentage of total lithospheric strength contributed by the crust is can be related to thermal perturbations of the sub-lithospheric upper highly variable covering the full spectrum from jelly-sandwich to mantle as imaged by seismic tomography (Goes et al., 2000), with later- crème brûlée end-member rheologies (Burov and Watts, 2006). al variations in crustal thickness playing a secondary role, apart from the Europe's lithosphere is characterized by major spatial mechanical Alpine domains which are characterized by deep crustal roots. The high strength variations, with a pronounced contrast between the strong strength of the East-European Platform, the Bohemian and London–

Fig. 9. Depth map of Moho discontinuity (Tesauro et al., 2008). Abbreviations are as follows: A, Apennines; AB, Alboran Basin; AP, Adriatic Promontory; BC, Betic Cordillera; BS, Black Sea; CH, Carpathians; CM, Cantabrian Mountain; D, Dinarides; EB, Edoras Bank; EL, ElbeLineament; EEP, East European Platform; FB, Focsani Basin, FI, Faeroe Islands; GB, Gulf of Bothnia; HB, Hatton Bank; IAP, Iberian Abyssal Plain; IS, Iapetus Suture; LVM, Lofoten–Vesteralen margin; MC, Massif Central; NGB, North German Basin; NS, North Sea; OR, OsloRift; P, Pyrenees; PB, Pannonian Basin; TS, Tyrrhenian Sea; TTZ, Tesseyre–Tornquist zone; URG, Upper Rhine Graben; VB, Vøring Basin; and VT, Valencia Trough.

Brabant Massifs and the Fennoscandian Shield reflects that they are have revealed the importance of its neotectonic structural reactivation underlain by old, cold and thick lithosphere, whereas the European Ce- by lithospheric folding (Marotta et al., 2000). Similarly, the Plio- nozoic Rift System coincides with a major axis of thermally weakened Pleistocene subsidence acceleration of the North Sea has been attributed lithosphere. Similarly, weakening of the lithosphere of southern France to folding. Lithospheric folding is a very effective mechanism for the can be attributed to the presence of tomographically imaged plumes propagation of tectonic deformation from active plate boundaries far rising up under the Massif Central (Granet et al., 1995; Wilson and into intraplate domains (e.g., Burov et al., 1993; Stephenson and Patterson, 2001). Furthermore, the strong Adriatic indenter contrasts Cloetingh, 1991; Ziegler et al., 1995, 1998, 2002). with the weak lithosphere of the Mediterranean collision zone. At the scale of a microcontinent that was affected by a succession The major lateral strength variations that characterize the litho- of collisional events, Iberia provides a well-documented natural labo- sphere of extra-Alpine Phanerozoic Europe are largely related to its ratory for lithospheric folding and the quantification of the interplay Late Cenozoic thermal perturbation as well as to Mesozoic and Cenozoic between neotectonics and surface processes (Cloetingh et al., 2002; rift systems and intervening stable blocks, and not so much to the Cale- Fernandez-Lozano et al., 2011, 2012). An important factor supporting donian and Variscan orogens and their accreted terranes (Dèzes et al., a lithosphere-folding scenario for Iberia is the compatibility of the 2004; Ziegler and Dèzes, 2006). These lithospheric strength variations thermo-tectonic age of its lithosphere and the wavelength of the (Fig. 10) are primarily the result of variations in the thermal structure observed deformations. of the lithosphere and, therefore, are compatible with inferred EET var- Well-documented examples of continental lithospheric folding also iations (see Cloetingh and Burov, 1996; Perez-Gussinye and Watts, come from other cratonic areas. A prominent example of lithospheric 2005). folding occurs in the Western Goby area of Central Asia, involving a Crustal seismicity is largely concentrated on the presently still active lithosphere with a thermo-tectonic age of 400 Ma. In this area, mantle Alpine plate boundaries, and particularly on the margins of the Adriatic and crustal wavelengths are 360 km and 50 km, respectively, with a indenter. In the Alpine foreland, seismicity is largely concentrated on shortening rate of 10 mm yr−1 and a total amount of shortening of zones of low lithospheric strength, such as the European Cenozoic 200–250 km during 10–15 My (Burov and Molnar, 1998; Burov et al., rift system, and areas where preexisting crustal discontinuities are 1993). Quaternary folding of the Variscan lithosphere in the area of reactivated under the presently prevailing NW–SE directed stress the Armorican Massif (Bonnet et al., 2000) resulted in the development field, such as the South Armorican shear zone and the fracture systems of folds with a wavelength of 250 km, pointing to a mantle-lithospheric of the Bohemian Massif (Dèzes et al., 2004; Ziegler and Dèzes, 2006), control on deformation. As the timing and spatial pattern of uplift and the rifted margin of Norway (Mosar, 2003). inferred from river incision studies in Brittany is incompatible with a glacio-eustatic origin, Bonnet et al. (2000) relate the observed vertical 5.3. Lithospheric folding: an important mode of compressional motions to deflection of the lithosphere under the present-day NW–SE reactivation of rifts directed compressional intraplate stress field of NW Europe. Stress- induced uplift of the area appears to control fluvial incision rates and Folding of the lithosphere, involving its positive as well as negative the position of the main drainage divides. This area, which is located at deflection (see Fig. 11), appears to play a more important role in the the western margin of the Paris Basin and along the rifted Atlantic mar- large-scale neotectonic deformation of Europe's intraplate domain gin of France, has been subjected to thermal rejuvenation during Meso- than hitherto realized (Cloetingh et al., 1999). The large wavelength of zoic extension related to North Atlantic rifting (Robin et al., 2003; Ziegler vertical motions associated with lithospheric folding necessitates inte- and Dèzes, 2006) and subsequent compressional intraplate deformation gration of available data from relatively large areas, often going beyond (Ziegler et al., 1995) that affected also the Paris Basin (Bourgeois et al., the scope of regional structural and geophysical studies that target 2007; Lefort and Agarwal, 2002,seeFig. 2B). Levelling studies in the specific structural provinces. Recent studies on the North German Basin Bretagne area (Lenotre et al., 1999) also point towards its ongoing

Fig. 10. Spatial variations in integrated strength and effective elastic thickness of Europe's lithosphere inferred from forward rheological modelling (from Tesauro et al., 2009).

Top) Percentage of total lithospheric strength due to the crust; Bottom) Effective Elastic thickness (Te) distribution of the European lithosphere derived from integrated strength of the lithosphere (km). Abbreviations are as follows: F, Fennoscandia; S, Sarmatia, SN, SvecoNorvegian; BS, Black Sea, M, Moesian Platform; C, Caledonides; V, Variscides; AD, Alpine Do- main; AM, Atlantic Margin; MS, Mediterranean Sea; and TESZ, Trans European Suture Zone. deformation. The inferred wavelengths of these neotectonic lithosphere of the lithospheric mantle is consistent with higher lithosphere growth folds are consistent with the general relationship that was established rates of lithosphere folding in the European foreland as compared to between the wavelength of lithospheric folds and the thermo-tectonic folding in Central Asia (Nikishin et al., 1993), which is characterized age of the lithosphere on the base of a global inventory of lithospheric by pronounced mantle strength (Cloetingh et al., 1999). folds (Fig. 12)(seealsoCloetingh and Burov, 1996, 2011). In a number of other areas of continental lithosphere folding, smaller wavelength 6. Models for continental breakup and rift basins crustal folds have also been detected, for example, in Central Asia (Burov et al., 1993; Nikishin et al., 1993). 6.1. Extensional basin migration: observations and thermo-mechanical Thermal thinning of the mantle–lithosphere, often associated with models volcanism and doming, enhances lithospheric folding and appears to control the wavelength and amplitude of folds, as seen by the Neogene Several examples of extensional basins have been described in uplift of the Rhenish Shield and the Vosges–Black Forest arch (Dèzes which the locus of extension shifted in time toward the zone of future et al., 2004; Ziegler and Dèzes, 2007). Substantial thermal weakening crustal separation (e.g., Bukovics and Ziegler, 1985; Lundin and Doré,

ocean boundary that was active during the stretching event preceding continental breakup at the Paleocene–Eocene transition (Lundin and Doré, 1997; Mosar et al., 2002; Osmundsen et al., 2002; Reemst and Cloetingh, 2000; Skogseid et al., 2000). The Nordland Ridge, forming the western margin of the Trøndelag Platform, is a classical footwall uplift that is associated with the border fault of the Vøring Basin, whilst the Vøring Marginal High is associated with the continent–ocean boundary. Although there is still considerable uncertainty about the age of the oldest syn-rift sedimentary sequences involved in the different deep-seated fault blocks of the Mid-Norway margin and the width of the area that was affected by the pre-Jurassic extensional phases (Gabrielsen et al., 1999; Mosar et al., 2002), it is generally agreed that in time extension shifted westward, away from the Trøndelag Platform

Fig. 11. Schematic diagram illustrating decoupled lithospheric mantle and crustal folding, towards the Vøring Basin and ultimately centred on the continental and consequences of vertical motions and sedimentation at the Earth's surface. V is horizontal breakup axis (e.g., Bukovics and Ziegler, 1985; Lundin and Doré, 1997; shortening velocity; upper crust, lower crust, and mantle layers are defined by corresponding Reemst and Cloetingh, 2000). Other areas in which extensional strain rheologies and physical properties. A typical brittle–ductile strength profile (in black) for concentrated in time toward the rift axis or the future breakup axis – – – decoupled crust and upper mantle lithosphere, adopting a quartz diorite olivine rheology, are, for instance, the Viking Graben of the North Sea (Ziegler, 1990), is shown for reference. the nonvolcanic Galicia margin and the passively rifted South Alpine Tethys margin (e.g. Bertotti et al., 1997; Manatschal and Bernoulli, 1997; Ziegler, 1988, 1996b). These basins typically consist of several 1999). laterally adjacent fault-controlled sub-basins, the main sedimentary Several hypotheses have been proposed in an effort to explain this fill of which often differs by several tens of millions of years. As such, lateral migration of rifting activity by invoking the principle that exten- they reflect lateral changes in their main subsidence phases that can sional strain is centred on the weakest part of the lithosphere (Steckler be attributed to a temporal lateral shift of the centres of lithosphere and Ten Brink, 1986). A mechanism for limiting extension at a given extension. A well-documented example of this phenomenon is the location was studied, for example, by England (1983), Houseman and 200 to 500 km wide Mid-Norway passive continental margin. Prior to England (1986),andSonder and England (1989), who found that the final Late Cretaceous–Early Tertiary rifting event that was accompa- cooling of the continental lithosphere during stretching may increase nied by the impingement of the Iceland Plume and culminated in conti- its strength, so that deformation shifts to a previously low strain region nental breakup, the Mid-Norway Vøring margin was affected by several (Sonder and England, 1989). With this mechanism, other effects such as more or less discrete rifting events (Bukovics and Ziegler, 1985; changes in plate boundary forces are not required to explain basin Skogseid et al., 1992; Ziegler, 1988). These resulted in the subsidence migration. of several depocenters located between the Norwegian coast and the Migration of the extension locus has also been attributed to tempo- continent–ocean boundary (Fig. 13B). On a basin-wide scale, this rally spaced multiple stretching phases, separated by periods during passive margin consists of an initial break-away zone of probably Late which the lithosphere is not subjected to extension but cools. Under Palaeozoic age located up to 150 km to the east of the coast, the Late such a scenario, the lithosphere that was weakened during an earlier Palaeozoic–Triassic Trøndelag Platform, located adjacent to the coast, stretching phase needs sufficient time to regain its strength by cooling the Jurassic–Early Cretaceous Vøring Basin in the central part of the and to become indeed stronger than the adjacent unextended area shelf, and the marginal extended zone adjacent to the continent– before the onset of the next stretching event. Obviously, this concept

Fig. 12. Theoretically predicted wavelengths as function of thermo-tectonic age (for different lithospheric layers as well as whole-lithosphere folding). Model is compared to the observed folding wavelengths (after Cloetingh et al., 1999). Thermo-tectonic age corresponds to the time elapsed since the last major perturbation of the lithosphere prior to folding.

Fig. 13. A) Tectonic subsidence curves for four sub-basins and platforms of the Mid-Norwegian margin. Note the simultaneously occurring subsidence accelerations (coloured bars). The Mesozoic subsidence accelerations reflect pronounced rift-related tectonic phases, whereas the Late Neogene subsidence acceleration (light yellow bar) coincides with a major reorganization of the Northern Atlantic stress field. Modified from Reemst (1995); B) Sketch map of rift zones on the Mid-Norwegian margin showing timing of main rift activity. Modified from Skogseid and Eldholm (1995).

Sawyer and Harry (1991) and Favre and Stampfli (1992) described plate separation is achieved. The duration of the rifting stage preceding for the evolution of the Central Atlantic rift a gradual concentration plate separation depends on the extension velocity. When the litho- of extensional strain towards the future zone of crustal separation. sphere is stretched at greater velocities, it takes less time to reach con- This phenomenon is attributed to the syn-rift gradual rise of the litho- tinental breakup. When extension velocities are less than 8 mm yr−1, spheric isotherms and a commensurate upward shift of the lithospheric stretching of the lithosphere does not lead to plate separation. The de- necking level and the intracrustal brittle/ductile deformation boundary. pendence of rift duration on potential mantle temperature is discussed As a result of this, extensional strain concentrates in time on the ther- in Van Wijk et al. (2001). The configuration of the rifted basins shows no mally more intensely weakened part of the lithosphere, generally clear dependence on the tested stretching velocities (see also Bassi, corresponding to the rift axis, thus causing narrowing of the rift and 1995). abandonment of its lateral graben systems (Ziegler and Cloetingh, The tendency for the lithosphere to neck (or to focus strain) is 2004). In the case of the Central Atlantic, this process was enhanced weaker with decreasing extension velocities. When stresses exerted by the impingement of a short-lived, though major, mantle plume at on the lithosphere are smaller, the rate of strain localization also the Triassic–Jurassic transition (Nikishin et al., 2002; Wilson, 1997). slows down, mantle upwelling is slower and syn-rift lateral conductive In the following, we discuss the implications of a viscoelastic plastic cooling plays a more important role. When the lithosphere is stretched finite element model for extension of a lithosphere with an initially at high rates, upwelling of mantle material is fast (almost adiabatic) symmetric upper-mantle weakness (for details see Van Wijk and with little or no horizontal heat conduction. The fast rise of hot astheno- Cloetingh, 2002). When large extension rates are applied, deformation spheric material further reduces the strength of the lithosphere in the becomes focused, causing lithospheric necking and eventually plate central region, with the consequence that lithospheric deformation separation. Hereafter, this is referred to as ‘standard’ rifting. A different and thinning accelerates even further. The result is a short rifting evolution of deformation localization takes place when the lithosphere stage preceding plate separation. is extended at low strain rates. In this case, the necking area may start to These model-derived conclusions on the duration of the rifting migrate laterally, and hence delay or even prevent plate separation. stage preceding plate separation need to be viewed in the context The modelled domain consists of an upper and a lower crust and a of natural examples as summarized in Fig. 5 (Ziegler, 1996b; Ziegler lithospheric mantle (Fig. 14) to which the rheological parameters of and Cloetingh, 2004; Ziegler et al., 2001). granite, diabase and olivine, respectively, have been assigned (Carter and Tsenn, 1987). In order to facilitate deformation localization, the 6.3. Thermo-mechanical evolution and tectonic subsidence during slow crust was thickened by 2 km in the centre of the domain, forming a extension linear feature parallel to the future rift axis. This causes localized weakening of the upper mantle and a corre- The lithosphere reacts differently to low and high extension sponding strength reduction of the lithosphere. A lithosphere with rates (Burov, 2007; Huismans and Beaumont, 2003; Van Wijk and this configuration is thought to be typical for orogenic belts after Cloetingh, 2002). Results of a representative model simulating an ex- post-orogenic thermal equilibration of their thickened crust (Henk, tension rate of 6 mm/yr over a period of 100 My are here presented 1999). The Mid-Norway margin is superimposed on the Caledonian to illustrate the thermal evolution of the lithosphere (Figs. 14 and 15). orogen, the crust of which was presumably still thickened prior to The total strength of the lithosphere, obtained by integrating the stress the Late Carboniferous onset of rifting (Skogseid et al., 1992). field over the thickness of the lithosphere (Ranalli, 1995), is also shown in Fig. 16.Duringthefirst 30 My after the onset of stretching, deforma- 6.2. Fast rifting and plate separation tion localizes in the centre of the model domain where the lithosphere was pre-weakened (Figs. 14–16). The top of the crust subsided, a sedi- As the architecture of rifts that developed in response to variable mentary basin developed and mantle material welled up. Subsequently, extension rates (‘standard necking cases’) does not differ significantly, as lithospheric stretching proceeds, temperatures begin to decrease a representative simulation of (model for) lithosphere stretching at a (see time slices corresponding to 45 My, 50 and 60 My, Fig. 15), in con- velocity of about 16 mm yr−1 is presented here. At the onset of litho- trast to what happened in the standard necking case shown in Fig. 14. spheric extension deformation is localized at the centre of the domain Development of the mantle upwelling zone ceases and the lithosphere on which the initial mantle weakness was imposed. Thinning of the cools in the centre of the domain. Cooling of the central zone continues crust and mantle lithosphere concentrates here, and mantle material while after 70 My increasing temperatures are evident on both sides starts to well up (Fig. 14B). Under persisting extension thinning of the of the previously extended central zone. As these new upwelling crust and mantle lithosphere continues, resulting after 27 My in plate zones develop further (Fig. 15, 110 My panel) two new basins develop separation. Continental breakup is here defined as occurring when the adjacent to the initial stage basin. Temperatures in the lithosphere are crust is thinned by a factor 20, though another factor or another defini- now lower below the initial basin as compared to the surrounding tion for plate separation could have been chosen. This definition corre- adjacent lithosphere; a ‘cold spot’ is present in an area that underwent sponds to about 40–50% of extension of the model domain at plate initial extension (see also corresponding lithosphere strength evolution separation. Thinning factors for the crust (δ) and lithospheric mantle in Fig. 16). This thermal structure is reflected by the surface heat flow, (β)areshowninFig. 14C and are defined as the ratio between the initial which is the surface heat flow values are lower in the first stage basin and the present thickness of the crust or lithospheric mantle, respec- than in the adjacent areas at 110 My along cross-section normal to the tively. The base of the lithospheric mantle is the 1300 °C isotherm. rift axis. Burov (2007) suggested that slow extension may be associated The centre of the model domain, on which the initial mantle with the development of gravitational instabilities in the lower litho- weakness was imposed, is the weakest part; this continues to be so spheric mantle along the margins of the area of mantle lithosphere until plate separation. Integrated strength values of the continental lith- thinning, resulting in its sinking into the asthenosphere, and, hence, in osphere vary between 1012 and 1013 N/m with the higher values char- “gravitational” thinning of the lithosphere by mantle lithosphere acterizing Precambrian shields (Ranalli, 1995). In our model the delamination (Fig. 17). These instabilities develop because the mantle strength of the lithosphere falls within this range. During rifting, the lithosphere is colder and thus denser than the upwelling hot astheno- strength of the lithosphere decreases with time, owing to its thinning sphere (1330 °C) while at each moment of time the degree of assimila- and heating as a consequence of its stretching and nonlinear rheology. tion of the mantle lithosphere by the asthenosphere and its density In other cases with higher constant extension rates the localization contrast with the later is limited by the ratio of the advection to diffu- of deformation is comparable to that discussed above. In all cases, defor- sion rate (Peclet number). If the extension rate is low enough, the char- mation concentrates on one zone in which thinning continues until acteristic time scale of growth of basal mantle–lithosphere gravitational

Fig. 14. Modelled thermal evolution of the lithosphere and evolution of thinning factors for a stretching rate of Vext = 16 mm/yr. Modified from Van Wijk and Cloetingh (2002). A) Initial model configuration; B) Thermal evolution at 2, 10, and 20 My after the onset of stretching; Note the changing horizontal and vertical scales in the panels indicating the changing sizes of the model domain upon stretching; C) Evolution of thinning factors of crust (δ) and mantle lithosphere (β) for Vext = 16 mm/yr. Breakup occurs after 27 My. Width of the model domain (horizontal axis) versus time (vertical axis). The width of the model domain increases as the lithosphere is extended. Modified from Van Wijk and Cloetingh (2002).

instabilities becomes comparable or even smaller than the characteris- where μeff is effective viscosity. Assuming a ductile flow law: tic time scale of the syn-rift phase. These instabilities may eventually ¼ : μ ðÞ= =ðÞðÞΔρ þ αρ Δ grow even at higher rates than the rate of other processes. Der 13 04 0 exp H nRT ux c 0 T gd L Hence the Rayleigh –Taylor instabilities may interplay with rifting d(1 − n)/n ⁎ − 1/n d 1/2 processes, resulting in asymmetric rifting and/or in increased amount with μ0 = e II (A ) , e II =(InvII (eij)) is the effective strain of thinning, specifically in the mantle part of the lithosphere. Mantle rate and A⁎ =1/2A·3(n + 1)/2 is the material constant, H is the activa- lithosphere delamination may also eventually lead to additional partial tion enthalpy, R is the gas constant, and n is the power law exponent. melting resulting from the ascent of the replacing asthenospheric mate- L is the final width of the rift and ux is the extension rate, Δρc + b rial to sub-Moho depths ( 50 km). It can be further suggested that such αρ0ΔT = Δρ, Δρc is the compositional density contrast, α is the delamination-caused partial melting can be generated even under coefficient of thermal expansion, ρ0 is the density at reference temper- non-volcanic margins. One of the other consequences of the mantle ature and ΔT is the temperature change in respect to the reference lithosphere instabilities is the development of series of periodic changes temperature. Burov (2007) has shown that gravitational instabilities in subsidence associated with delamination of dense mantle. To charac- play an important role for Der b 10, specifically for Der b 1.5. It is also terize the relative role of the gravitational instabilities during rifting, noteworthy that Moho depth and geometry are largely controlled by Burov (2007) introduced a special parameter, the “rift Deborah number”: extension and the evolution of mantle-lithospheric gravitational instabilities. In particular, slow extension may result in subsidence of the Moho during the syn-rift phase. Fig. 18 shows a critical case at an extension rate De ¼ 13:04 μ u =ðÞΔρgd L r eff x of 10 mm/yr corresponding to Der ~ 1, below which, inclusively, rifting

Fig. 15. A) Thermal evolution of the lithosphere for a migrating rift at an extension rate of Vext = 6 mm/yr, times in My after the onset of stretching.; B) Evolution of thinning factors for the crust (δ) and mantle (β) at an extension rate of Vext = 6 mm/yr. No continental breakup. Modified from Van Wijk and Cloetingh (2002). can be treated as “slow” in terms of the impact of gravitation), so that extension at rates exceeding the extension rate. Hence, for relatively Moho remains almost flat or goes down instead of rising up. Hence, in high extension rates it may occur at some distance from the rift flanks. the presence of a deep lithospheric necking level, and under low exten- It can be questioned if the impact of the gravitational instabilities sion rates, the Moho may remain stable or even subside in response to will be smaller in case of differential stretching of the lithosphere lower mantle-lithosphere gravitational instabilities. This does not neces- with mantle β-factors stronger than the crustal ones. β-Factors, how- sarily mean that upon detachment of significant volumes of lithospheric ever, are not connected to a particular mechanism of lithospheric mantle, the lithosphere rebounds and the graben flanks can be uplifted, thinning: the gravitational instabilities that erode mantle lithosphere even in the absence of further extension. The detachment may occur in in addition to the effect of pure or simple shear mechanism are by pace with thermal re-equilibration of the lithosphere so that the density themselves a cause of differential stretching. contrast between the latter and the underlying mantle can be quite small Thinning factors for the crust and mantle lithosphere corresponding at the moment when the detachment finally takes place. The models also to the experiments shown in Figs. 15–16 vary at a large extent. Thinning show that the detachment may laterally migrate in the direction of the of the crust starts, as expected, in the central weakness zone of the domain where a single basin forms, with a maximum thinning factor of 1.85 for the crust. Crustal thinning in the central basin continues until about 65 My after the onset of stretching, at which time the locus of thinning shifts towards both sides of the initial basin. The zones of second-stage maximum thinning are located at a distance of about 500 km from the centre of the initial rift. Although stretching of the lithosphere continues after 65 My, extensional strain is no longer localized in the initial rift basin but is centred on the two flanking new rifted basins. The thinning factor of the mantle lithosphere reflects this behaviour. During the first 45 My of stretching upwelling mantle material is primarily localized in the central zone rather than in its surroundings. Thereafter, however, temperatures decrease rapidly in the central zone (see also the associated lithosphere strength variations in Fig. 16 as new upwelling zones develop on its flanks with decreasing activity in the area of the first basin and increasing upwelling beneath the new basins. After the onset of stretching (Fig. 16) the central part of the model domain progressively weakens, reaching its minimum strength by 30 My. Thereafter its strength increases, but remains lower than the Fig. 16. Lithosphere strength evolution for a migrating rift, Vext = 6 mm/yr. Modified from Van Wijk and Cloetingh (2002). rest of the domain. From 55 My onward, the smallest values of

Fig. 17. Simplified cartoon showing interaction between the asthenosphere and mantle lithosphere during slow rifting (Burov, 2007). Left: The mantle lithosphere below the rift flanks becomes gravitationally unstable due to the negative density contrast with the hot asthenosphere and the loss of mechanical resistance resulting from lateral heat transfer from the asthenosphere below the rifted zone. The asthenosphere is positively unstable because of its positive density contrast with the embedding mantle lithosphere. Since the viscosity of the mantle lithosphere is exponentially dependent on temperature, there is a relatively net “stability level” that separates the regions of high viscosity from thermally weakened regions of low viscosity. These regions are subject to development of rapid RT instabilities. Right: simplified yield-stress pre-rift envelope of the continental lithosphere and typical pre-rift temperature profile. lithospheric strength occur on both sides of the central basin beneath lithosphere) and uplift of the rift shoulders, create pressure gradients the second-stage rifts. By comparing the strength of the lithosphere sufficient to drive ductile flow in the low-viscosity lower crust with its thermal structure (Fig. 16), the strong dependence of (Fig. 19, Burov and Cloetingh, 1997). Ductile flow occurs until the duc- strength on the temperature is evident. From this modelling study it tile channel thickness is greater than a few km (Burov and Cloetingh, is concluded that lateral rift migration may occur under conditions 1997). This flow is also largely gravity driven, due to lateral Moho gra- of long-term low-strain-rate lithospheric extension. Similar results dients associated with crustal thinning. Hence, the flow rate first pro- have been also obtained in the earlier study by Burov and Poliakov gressively increases as the crust thins, and then decreases when there (2001) who have demonstrated a potentially strong impact of litho- is no more ductile crustal material below the rift zone. This flow, direct- spheric strengthening due to mechanical coupling of rheological layers ed outward from the centre of the basin may facilitate uplift of the rift occurring at some degree of thinning, specifically when the rate of shoulders (Fig. 20, Burov and Poliakov, 2001). It may even drive some cooling of the lithosphere is compatible or higher than the advection post-rift “extension” (Burov and Poliakov, 2001). In the limiting case rate. Even if we are unaware of present-day examples of two younger of slow erosion and sedimentation rates, gravitational stresses can re- rifts flanking an older rift, the exposed mechanism could have played verse the flow, resulting in retardation of basin subsidence rate, homog- an important role in past rifting tectonics. Its viability can be demon- enisation of the crustal thickness, accelerated collapse of the shoulders strated, for example, on the case of the East-African rift, which is and in some post-rift “compression”. These effects may significantly subdivided by strong Tanzanian craton into two simultaneously local- change predictions of basin evolution inferred from the conventional ized rifts (Western and Eastern branches, e.g., Corti et al., 2007). stretching models. It is noteworthy, however, that it is rather difficult to refer to direct observations allowing for testing crustal flow models, as well as any models of subsurface dynamics, against natural cases. 6.4. Interplay of lower crustal flow and surface erosion in rifts Model-driven hypothesis can be verified indirectly by showing that simpler models fail to explain vital observations such as poly-phase Surface processes play a very important role during both syn-rift and subsidence, post-rift doming and flank uplift. post-rift evolution (Burov and Cloetingh, 1997; Burov and Poliakov, 2001), not only in terms of their effect of thermal blanketing causing temperatures to rise in the sedimentary cover, but also due to redistri- 6.5. Breakup processes: timing and mantle plumes bution of surface loads (Fig. 19). Surface processes modify topography and sedimentary rates comparable with the rate of the tectonic uplift/ Continental breakup is a consequence of plume–lithosphere interac- subsidence (a few 0.1 mm/yr), leading to the erosion of thousands tions when a low density, high temperature and low viscosity mantle meters of topography from uplifted rift flanks and deposition of equiv- plume, separated from large-scale convective motions, rises up from alent amount of sedimentary infill. The associated dynamic loading and the core–mantle boundary to the base of the lithosphere (e.g., Sleep, unloading exerted on the supporting crust and lithosphere are signifi- 2006) applying normal and basal shear stresses and producing cant, suggesting possible feedbacks between the surface and tectonic magmatism associated with decompressional partial melting and processes. In particular, an increase of sedimentary loading leads to thermo-mechanical erosion that eventually lead to thinning and rup- localised yielding and deflection of the supporting lithosphere. At the ture of the overlying plate. This upwelling, if continuously fed from a same time, erosional unloading of rift shoulders leads to their flexural deeper source area, forms a “superplume”, i.e. a persisting diapir of rebound and strengthening. These processes, resulting in acceleration very large scale coming from the D″ boundary (e.g. Condie et al., of the subsidence of the rift “neck” (strongest layer of the thinned 2000; Romanowicz and Gung, 2002). There may be also smaller scale

Fig. 18. Results of the experiments (Burov, 2007) for extension rate (10 mm/yr, Der 1) assuming compositional density contrast between mantle and asthenosphere, Δρc = 20 kg m-3. As suggested by the corresponding value of Der (1), 10 mm/yr is a rate below which the rifting can be treated as “slow” in terms of the possibility of the development of gravitational instabilities. Left: evolution of the material field (colour code: upper crust — salad-green; middle-lower crust — yellow; mantle lithosphere — blue; asthenosphere — marine-green; syn-rift sediment or otherwise reworked material — purple). Centre: temperature field. Left: logarithm of stress to strain ratio (Pa s), which is equivalent to the effective viscosity for the ductile domains. Note: 1) removal of a large part of the mantle lithosphere by RT instabilities leads to change of style of rifting; 2) Strong asymmetric rifting at developed stages of extension. 3) Breakup and oceanization of the lithosphere that commences at 9 Myr and is preceded by coupling of the crustal and mechanical layers within the lithosphere. 4) Exhuma- tion of small amounts of lower crust and mantle. Arrows show the position of the interior basins. Inserts correspond to 2× blow-ups of the framed areas. Purple (reworked) material most often directly overlies its source material, i.e. purple above green layer means material reworked from the upper crust. Purple above blue means either reworked mantle or sediment derived from the upper and lower crust. Note: 1) removal of a large part of the mantle lithosphere by RT instabilities leading to change of style of rifting; 2) strong asymmetric rifting at developed stages of extension (at the oceanization stage); and 3) exhumation of continental and oceanic mantle and punctual exhumation of lower crustalmaterialatocean–continent limits. In this experiment, continental break up and oceanization commence at 6 Myr. Arrows show the position of the interior basins. short-lived diapirs originating from different depths in the upper man- systems created at the edges of the African cratonic lithosphere (Corti, tle (Courtillot et al., 2003; Montelli et al., 2004; Ritter, 2005). The 2005; Janssen et al., 1995). It should be noted that some examples do resulting thermo-mechanical consequences for surface geodynamics not always fit these assumptions. For instance, the Labrador Sea opened and geology appear to be largely dependent on not only plume charac- orthogonally to the basement grain, whereas the North Sea rift cuts teristics, but also on the thermo-rheological and density structure of the obliquely across the Caledonian basement grain (Ziegler, 1988). lithosphere (Burov and Cloetingh, 2010; Burov and Guillou-Frottier, Plume impingement may also occur at the onset of the post-rift phase 2005; Burov et al., 2007). Both the thermo-tectonic age and rheological or even later. In this case the rifted margin lithosphere has been ther- stratification of the lithosphere have a strong impact on the effect on mally reset to very young thermo-mechanical ages for the thinned plume impingement (Burov and Guillou-Frottier, 2005). As a result, continental lithosphere and the adjacent oceanic lithosphere. the presence of low-viscosity ductile lower crust may result in essential The interplay between extension and magmatism during continen- damping of the long-wavelength dynamic topography and appearance tal breakup is still debated and recent numerical modelling studies of short-wavelength tectonic scale features, associated with tensional suggest that the volumes of melts extruded at volcanic margins may and compressional instabilities in the crustal layers. The assumption also be generated by ‘standard’ thermal conditions, provided high ex- of a horizontally uniform lithosphere at the site of future break-up is tension rates can be implied (Huismans and Beaumont, 2011; Van also not realistic, in view of the abundant evidence from the geological Wijk et al., 2001). record that incipient rifts and rifted margins are usually localized at Volcanic margins such as the Mid-Norway margin are a representa- suture zones separating stronger lithosphere. Examples include the tive part of the Norwegian–Greenland Sea rift along which crustal Caledonides suture of Laurentia–Greenland and Baltica, localizing Late separation between NW Europe and Greenland was achieved in earliest Carboniferous and Permo-Triassic rifting and subsequent continental Eocene times (Mosar et al., 2002; Torsvik et al., 2001). The Norwegian– break-up around 65 Ma in the Arctic–Northern Atlantic and the rift Greenland Sea rift, which had remained intermittently active for some

Fig. 19. Proposed scenario of basin subsidence: sediments derived from erosion on slopes of rift shoulders result in increase of loading in the basin. Strong parts of the crust and of mantle lithosphere bend and undergo flexural weakening at the inflexion points (Burov and Cloetingh, 1997). As result, the equivalent elastic thickness, EET, drops beneath the basin and rift shoulders (bottom) and becomes lower than the EET immediately after extension. Lower crustal material flows from the centre of basin towards the shoulders facilitating their uplift.

Fig. 20. Numerical experiments showing impact of surface processes on syn-rift evolution (modiﬁed from Burov and Poliakov, 2001). Shown are two cases, with and without erosion. Thermo-tectonic age of the lithosphere in this experiment is 400 Ma, the rheological envelope includes 3 layers: 2 layer crust (quartz-dominated upper crust and quartz–diorite lower crust) and dry olivine mantle. Note that erosion accelerates rift ﬂank uplift and allows localization of deformation on major border faults.

Fig. 21. Results of gravity modelling for the Eastern Black Sea, demonstrating isostatic ﬂexural overcompensation in the centre of the basin. For location see Fig. 24, line C–C′. Modiﬁed from Cloetingh et al. (2003).

280 Ma from the Late Carboniferous until the end of the Palaeocene, to conceptual models (Fig. 4), the Moho of a rift zone is uplifted in case forms part of the Arctic–North Atlantic rift (Larsen et al., 1999; of a deep necking level (strong lithospheric mantle) or downwarped in Skogseid et al., 2000; Ziegler, 1988). Tomographic data image a mantle case of a shallow necking level (weak lithospheric mantle). It is note- plume beneath Iceland rising up from near the core–mantle boundary worthy that this is independent from plume-driven active or regional (Bijwaard et al., 1998; Rickers et al., 2013). It should be noted that a stress-driven passive rifting (Ziegler, 1992, 1994) but is related to the temporal relationship between plume activity and the duration of strength of the lithosphere that is subjected to extension. As mentioned rifting is lacking (e.g. Ziegler et al., 2001). in previous sections, in nature rifting might never occur in a purely active or passive mode. At least, the presence or absence of surface 6.6. The evolution of the Moho geometry during rifting expression of syn-or pre-rift volcanic activity below the rifted continental crust cannot be directly associated with a particular mechanism of There are only few data that allow to constrain the evolution of the rifting (e.g., Huismans and Beaumont, 2011). Extension that started in Moho boundary during the syn-rift phase (e.g., Figs. 21, 22). According a passive mode may result in convective instability in the underlying

Fig. 22. Crustal thickness in the Pannonian Basin and the surrounding mountains. Values are given in km. 1: foreland (molasse) foredeep; 2: ﬂysch nappes; 3: pre-Tertiary units on the surface; 4: Penninic windows; 5: Pieniny Klippen Belt; and 6: trend of abrupt change in crustal thickness. Modiﬁed from Horváth et al. (2006).

Fig. 23. Time evolution of the model from the experiments showed in Fig. 20. As can be seen, tensional stresses are ﬁrst concentrated in the mantle part, and at this initial stage the Moho goes down. Once the mantle layer is thinned, the deformation is migrating to the upper layers, and the Moho goes up.

6.7. Post-rift inversion, intraplate stresses, borderland uplift, and understanding of the relationship between changes in plate motions, denudation plate interaction, and the evolution of rifted basins (Doré et al., 1997; Janssen et al., 1995; Johnson et al., 2008) and foreland areas (Ziegler Mechanisms controlling the development of inversion structures et al., 1998, 2001). remain enigmatic as theoretical models predict the inversion of passive A continuous spectrum of stress-induced vertical motions can margins in response to the buildup of compressional ridge-push forces occur in the sedimentary record, varying from subtle faulting effects only a few tens of million years after crustal separation has been to basin inversion (Brun and Nalpas, 1996; Ter Voorde and Cloetingh, achieved. Most of the shortening on the Mid-Norway margin was 1996; Ter Voorde et al., 2007; Ziegler et al., 1998) to the enhance- accommodated along preexisting major fault zones (Gabrielsen et al., ment of flexural effects and to lithospheric folding induced by high 1999; Pascal and Gabrielsen, 2001). Compressional structures, such as levels of stress approaching lithospheric strengths (Bonnet et al., long-wavelength arches and domes, strongly modified the architecture 1998; Burov et al., 1993; Cloetingh and Burov, 2011; Cloetingh of the deep Cretaceous basins and controlled sedimentation patterns et al., 1999; Stephenson and Cloetingh, 1991). during Cenozoic (Bukovics and Ziegler, 1985; Vågnes et al., 1998). Numerical models have been developed for simulating the interplay From the Oligocene onward, the near-shore parts of the Mid- of faulting and folding during intraplate compressional deformation Norway margin were uplifted and deeply truncated (Doré and Jensen, (Beekman et al., 1996; Cloetingh et al., 1999; Gerbault et al., 1998). 1996). Mechanisms controlling the observed broad uplift of the inner Models have also been developed to investigate the effects of faulting shelf and the adjacent on-shore areas, as evident along the Norwegian on stress-induced intraplate deformation in rifted margin settings coast also remain enigmatic. However, the long-wavelength of the (Van Balen et al., 1998). uplifted area is suggestive for mantle processes (Rohrman and Van In general, compressional intraplate tectonics appears to have der Beek, 1996; Rohrman et al., 2002). Olesen et al. (2002) interpreted a strong impact on the post-rift evolution of extensional basins the long-wavelength component of the gravity field in terms of (e.g. transpressional fault reactivation causing uplift of anticlinal both Moho topography and large-scale intrabasement density structures, phases of accelerated basin-wide subsidence in response contrasts. to stress-induced deflection of the lithosphere (see for a recent over- The seaward facing part of uplifted land areas was affected by strong view, Cloetingh et al., 2008). glacial erosion, which in turn enhanced their uplift and increased sedimentation and subsidence rates in the flanking basins (Cloetingh et al., 7. Back-arc basin formation and evolution 2005b). Fission track geothermal chronology analyses along major on-shore lineaments of southern Norway show that preexisting major Back-arc extensional basins evolve in response to a decrease in normal faults dating back to the Late Palaeozoic and Mesozoic, played the convergence rate or even a temporary divergence of colliding asignificant role in the Cenozoic uplift pattern (Cloetingh et al., plates, causing steepening and roll-back of the subducting lower 2005b; Hendriks and Andriessen, 2002; Redfield et al., 2005). Under plate lithospheric slab and the development of a secondary upwelling the presently prevailing northwest-directed compressional stress field system in the upper plate mantle wedge above the subducting slab the Mid-Norwegian margin and its adjacent highlands are seismically (e.g. Doglioni et al., 2007; Honza, 1993; Tamaki and Honza, 1991; active (Grünthal, 1999) with some faults showing evidence for recent Uyeda and McCabe, 1983). Back-arc basins can be developed over movement (Mörner, 2004). Uplift of the South Norwegian highland oceanic as well as continental lithosphere. Continental back-arc rifting continues, whilst the North Sea Basin experiences a phase of accelerated can progress to the opening of limited oceanic basins (e.g. Sea of subsidence that began during the Pliocene and is attributed to Japan, Black Sea, South China Sea). However, as convergence rates stress-induced deflection of the lithosphere (Van Wees and Cloetingh, between colliding plates vary in time, back-arc extensional basins are 1996). generally short-lived. Upon a renewed increase in convergence rates Southwestern Norway was uplifted by as much as to 2 km during and swallowing of the subducting lower plate lithospheric slab, Neogene times (Rohrman et al., 1995), as evidenced by the progradation back-arc extensional basins are subjected to compressional stresses of clastic wedges into the North Sea Basin (Jordtetal.,1995). However, that can lead to their inversion and closure. Typical examples may the uplift patterns and timing of southwestern Norway and the northern include the Variscan geosyncline Sunda arc and East China or the Scandes differ (Hendriks and Andriessen, 2002). By Late Tertiary times, Black Sea (Uyeda and McCabe; Cloetingh et al., 1989; Hall et al., 2011; cold climatic conditions prevailed. Jolivet et al., 1994; Munteanu et al., 2011; Nikishin et al., 2011; Intraplate stresses can significantly influence the post-rift evolution Ziegler, 1990). The sedimentary geometry and architecture of back-arc of extensional basins, particularly when transmitted from orogens into basins is controlled by such parameters as the age of subducted adjacent back-arc basins, such as the Black Sea or the Pannonian Basin lithosphere, the subduction direction, the type of underlying litho- (e.g., Bada et al., 2007; Cloetingh et al., 2003; Grenerczy et al., sphere (oceanic versus continental) or the uplift of the orogenic arc 2005). Field studies, kinematic indicators and numerical modelling (e.g., Dewey, 1981; Doglioni et al., 2007; Mathisen and Vondra, 1983; of present-day and palaeo-stress fields in selected areas (e.g., Bada Uyeda and Kanamori, 1979). These parameters control the large variety et al., 1998, 2001; Gölke and Coblentz, 1996) yielded new constraints of back-arc basins of various ages presently overlying different types of on the causes and the expression of intraplate stress fields in the lith- crust, such as the Caribbean, Banda-Sunda back-arcs or the Black Sea osphere. Ziegler et al. (2002, 2006) discussed the role of mechanical Basin (e.g., Hall et al., 2011; Meschede and Frisch, 1998; Munteanu controls on collision related compressional intraplate deformation. et al., 2011; Spakman and Hall, 2010). Temporal and spatial variations in the level and magnitude of these An observation that is common to many extensional back-arc basins stresses have a strong impact on the record of vertical motions in floored by continental crust is the discrepancy between the observed sedimentary basins (Cloetingh and Kooi, 1992b; Cloetingh et al., amounts of crustal and lithospheric mantle stretching, as derived from 1990; Zoback et al., 1993). Stresses propagating from the margins the magnitude the upper crustal faulting and crustal/lithospheric thick- of plates spreading axes into their interior had not only a strong nesses. This type of back-arc basin is often characterized by a much effect on their stratigraphic record, but can also contribute to ob- thicker post-rift basin fill than predicted by syn-rift stretching as seen served late-stage subsidence accelerations (Cloetingh and Burov, in the Pannonian Basin (Horváth et al., 2006; Lenkey, 1999). Such con- 2011; Cloetingh et al., 1999), similar to what is recognized in the trasting features occur in the hinterland of subduction driven orogens Pannonian Basin and the North Sea Basin (Horváth and Cloetingh, (sensu Doglioni et al., 2007; Royden and Burchfiel, 1989), as for exam- 1996; Van Wees and Cloetingh, 1996). Over the last few years, in- ple the case of subduction driving the highly arcuate Mediterranean creasing attention has been directed to this topic, advancing our orogens (e.g., the Pannonian–Transylvanian system, the Aegean Sea or

Fig. 24. Tectonic map of the Black Sea and adjacent areas (simplified and modified from Munteanu, 2012) with the location of profiles in Fig. 25. Continuous lines are faults separating tectonic units in and around the Black Sea. dashed lines are zones of diffuse and large scale deformation. NAF — North Anatolian Fault; WCF — West Crimea Fault; OF — Odessa Fault; BF — Bistrita Fault; TF — Trotus Fault;.

Fig. 25. Crustal scale models for basin evolution for the Western and Eastern Black Sea. For location of cross sections see Fig. 24. A comparison of predicted and observed Moho depths provides constraints on levels of necking and thermal regime of the pre-rift lithosphere. The models support the presence of cold pre-rift lithosphere compatible with a deep necking level of 25 km in the Western Black Sea. For the Eastern Black Sea, the models suggest the presence of a warm pre-rift lithosphere with a necking level of 15 km. Modified from Cloetingh et al. (2003). eastern Black Sea Basin. These findings raise important questions on of rift shoulders in the western Black Sea Basin as compared to the east- post-rift tectonics and on intraplate stress transmission into the ern Black Sea (Robinson et al., 1995). Black Sea Basin from its margins. 7.1.2. Strength evolution and neotectonic reactivation at the basin 7.1.1. Rheology and back-arc rift basin formation margins during the post-rift phase Inferred differences in the mode of basin formation between the Automated backstripping analyses and comparison of results with western and eastern Black Sea (Figs. 24–27), basins can be largely forward models of lithospheric stretching (Van Wees et al., 1998) explained in terms of paleo-rheologies. The pre-rift lithospheric provide estimates of the integrated lithospheric strength at various strength of the western Black Sea (Fig. 25) appears to be primarily syn- and post-rift stages. controlled by the combined mechanical response of a strong upper Fig. 26 shows a comparison of observed and forward modelled crust and strong upper mantle (Spadini et al., 1996). The shallow tectonic subsidence curves for the centre of the Western Black necking level in the eastern Black Sea is compatible with a pre-rift Sea Basin. Automated backstripping yields a stretching factor of 6. strength controlled by a strong upper crust decoupled from the Modelling, however, fails to predict the pronounced Late Neogene sub- weak hot underlying mantle (Fig. 25). These contrasts point to impor- sidence acceleration, documented by the stratigraphic record that may tant differences in the thermo-tectonic age of the lithosphere of the be attributed to late stage compression. Because post-rift cooling of two sub-basins (Cloetingh and Burov, 1996). The inferred lateral varia- the lithosphere leads in time to a significant increase in its integrated tions between the western and eastern Black Sea suggest thermal stabi- strength, its early post-rift deformation is favoured. Present-day litho- lization of the western Black Sea prior to rifting whilst the lithosphere of spheric strength profiles calculated for the centre and margin of west- the eastern Black Sea was apparently already thinned and thermally ern Black Sea show a pronounced difference. On the contrary, the destabilized by the time of Eocene rifting. The inferred differences onset of the basin inversion during late Middle Eocene times at short in necking level and in the timing of rifting between the western and times after the extension ceased earlier during Middle Eocene means eastern Black Sea suggest an earlier and more pronounced development that the lithosphere was weak when compression was activated. The

Fig. 26. Upper panel — Comparison of observed and forward modelled tectonic subsi- Fig. 27. Comparison of observed and forward modelled tectonic subsidence for the dence for the Western Black Sea centre. Automated backstripping yields an estimated Eastern Black Sea centre. Automated backstripping yields an estimated stretching factor stretching factor of 6 (top panel). Post-rift cooling leads in time to a significant increase of 2.3 (upper panel). Post-rift cooling leads in time to a significant increase in the predicted in the predicted integrated strength for both compressional and extensional regimes integrated strength for both compressional and extensional regimes (middle panel 1 (middle panel; 1 TN/m = 1012 N/m); Middle panel — Present-day lithospheric compres- TN/m = 1012 N/m). Present-day lithospheric strength profiles calculated for the sional strength profiles calculated for the centre and margin of western Black Sea show centre and margin of the Eastern Black Sea show with depth a pronounced difference with depth a pronounced difference (bottom panels). Temperature profiles (in C) and (bottom panels). Moho depth are given for reference. Note the important role of the actual Moho position Modified from Cloetingh et al. (2003). in terms of mechanical decoupling of the upper crust and mantle parts of the Black Sea lithosphere. Modified from Cloetingh et al. (2003). Horizontal lines indicate corresponding strength of the lithosphere at the time of the reversal of the stress field from the initial Cretaceous–Middle Eocene extension (lower horizontal line) to late Middle Eocene a cooling-related progressive increase of the integrated lithospheric compression (upper horizontal line) in the Black Sea Basin. The latter marks the onset of strength, with time increasingly higher stress levels are required to the inversion in the basin. cause large-scale deformation (Figs. 26 and 27). This is probably why the amplitude of contractional deformation decreases in time in the presence of relatively strong lithosphere in the basin centre in the western Black Sea basin (Munteanu et al., 2012). Important in this con- very late stages of Pliocene inversion and weaker lithosphere at the text is also that, due to substantial crustal thinning, a relatively strong basin margins favours deformation of the latter during late-stage upper mantle layer is present in the central part of the basin at relatively compression. shallow depths that may contrast substantially with an inherited weak- In Fig. 27 the observed and forward modelled tectonic subsidence er lower crust and, therefore, favour the transmission of contractional curves for the centre of eastern Black Sea Basin are compared, adopting stresses over large distances (Munteanu, 2012). for modelling a stretching factor of 2.3 that is compatible with the sub- Based on the present thermo-mechanical configuration (Figs. 26 sidence data and consistent with geophysical constraints. During the and 27) with relatively strong lithosphere in the basin centre and rel- first 10 My of post-rift evolution, integrated strengths are low but sub- atively weak lithosphere at the basin margins, we predict that the bulk sequently increase rapidly owing to cooling of the lithosphere. This of shortening will be taken up in the centre of the basin during the initial would correspond to the initial moments of post-Middle Eocene inver- phase of shortening, while the late-stage shortening induced by orogen- sion. A short time after the ultimate cessation of rifting (~10 My) in ic activity in the surrounding areas will be taken up along the basin Middle Eocene times, the basing was still underlain by very weak and margins, with only minor deformation occurring in the relatively stiff extended lithosphere. This is the moment when the entire Black Sea by now central parts of the basin. The relative difference in rheological basin was inverted due to the collision recorded in the Pontides domain strength of the marginal and central parts of the basin is more (e.g., Munteanu et al., 2011). It should be noted, however, that with pronounced in the eastern than in the western Black Sea. These

Fig. 28. Tectonic map of the Alps–Carpathians–Dinaridic system (modiﬁed after Matenco and Radivojević, 2012; Schmid et al., 2008) with the extent of the large Pannonian and Transylvanian back-arc basins and the location of other Miocene basins superposed over the Dinarides and Carpathians structures. 1 is the location of the cross section presented in Fig. 31. Red dashed oval is the location of the epicentres of the Vrancea seismogenic zone in the SE Carpathians.

Please cite this article as: Cloetingh, S., et al., The Moho in extensional tectonic settings: Insights from thermo-mechanical models, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.06.010 34 S. Cloetingh et al. / Tectonophysics xxx (2013) xxx–xxx stage for numerical and analogue modelling (e.g., Bada et al., 1999; in the Pannonian Basin, suggesting a two-phase evolutionary scheme Dombrádi et al., 2010; Jarosinski et al., 2011). for the system. According to their model, the initial early Middle The Pannonian–Carpathian system, therefore, permits to test Miocene basin subsidence was driven mainly by passive rifting in models for the development of sedimentary basins and their subse- response to roll-back of the subducted Carpathian slab, involving quent deformation, and for on-going continental collision. The litho- gravitational collapse of the thickened pre-rift Pannonian lithosphere. sphere of the Pannonian Basin is a particularly sensitive recorder of This triggered small-scale convective upwelling of the asthenosphere changes in lithospheric stress induced by near-field intraplate and that has favoured the late-stage thermal subsidence (Horváth, 1993) far-field plate boundary processes (Bada et al., 2001). as well as the compressional inversion of the Pannonian domain during Late Miocene–Pliocene times (e.g., Bada et al., 1999; Fodor 7.2.1. Neogene development and evolution of the Pannonian Basin et al., 2005). The Pannonian Basin of Central Europe (Fig. 28) is a classical exten- One other category of models involves lithospheric mantle insta- sional back-arc basin formed during Miocene times in response to the bilities for the formation of the asthenospheric upraise beneath the rapid roll-back of a slab attached to the European continent, which is Pannonian basin and the high-velocity anomalies observed by still underlain by highly thinned continental lithosphere (e.g., Balla, teleseismic tomography in neighbouring orogens (e.g., Dando et al., 1986; Cloetingh et al., 2006; Horváth, 1993; Horváth et al., 2006; 2011; Gemmer and Houseman, 2007; Houseman and Gemmer, Royden, 1988). The roll-back is partly responsible for the creation of 2007; Ren et al., 2012). These models assume a convective mantle re- the highly arcuate geometry of the Carpathian Mountains (e.g., moval beneath the Pannonian basin, correlated with high heat flow Matenco et al., 2010), a process that is common with many other Med- and interpreted as upwelling asthenosphere correlated with a fast iterranean orogens (e.g., Faccenna et al., 2004; Jolivet and Faccenna, anomaly extending outwards as far as the Carpathians, the Dinarides 2000). The Pannonian basin is spatially juxtaposed over an orogenic and the Eastern Alps. This higher velocity mantle material is area that formed in Cretaceous–Paleogene times in response to the sub- interpreted as being produced by a mantle downwelling, whose de- duction and closure of a number of oceanic domains which were kine- tachment from the lithosphere above may have triggered the exten- matically linked with the evolution of the larger Alpine Tethys and sion of the Pannonian Basin. Neotethys oceans (e.g., Csontos and Vörös, 2004; Schmid et al., 2008). Evolutionary models of the Pannonian Basin assume the onset of exten- 7.2.3. Stretching models and subsidence analysis sion at ~20 Ma, subsequently followed by a peak tectonic activity Sclater et al. (1980) were the first to apply the stretching model of along normal faults during Middle Miocene times, which was subse- McKenzie (1978) to the intra- Carpathian basins. They found that the quently followed by a post-rift, thermal sag phase starting in late development of peripheral basins could be fairly well simulated by Miocene times (e.g., Tari et al., 1999 and references therein). A latest the uniform pure shear extension concept with a stretching factor of Miocene–Quaternary contractional event has subsequently overprinted about 2. For the more centrally located basins, however, their consider- the basin during the translation and counter-clockwise rotation of able thermal subsidence and high heat flow suggested unrealistically the Adriatic indenter (Bada et al., 2007; Fodor et al., 2005; Horváth, high stretching factors up to 5. Thus, they postulated differential exten- 1995; Horváth and Cloetingh, 1996; Pinter et al., 2005). sion of the Pannonian lithosphere with moderate crustal stretching (δ factor) and larger stretching of the lithospheric mantle (β factor). 7.2.2. Dynamic models of basin formation Building on this and using a wealth of well data, Royden et al. (1983) Several models have been proposed to explain the dynamics of Neo- introduced the non-uniform stretching concept according to which gene rifting in the Pannonian Basin. An active versus passive mode of the magnitude of lithospheric thinning is depth-dependent. This rifting has been a matter of continued debate (see Bada and Horváth, concept accounts for a combination of uniform mechanical extension 2001) resulting in the proposal of various dynamic models that take of the lithosphere and thermal attenuation of the lithospheric mantle into account such prominent features of the Pannonian–Carpathian (Ziegler, 1992, 1996b; Ziegler and Cloetingh, 2004). This is compatible system as thinned and hot versus thickened and colder lithosphere in with the subsidence pattern and thermal history of major parts of the its central and peripheral sectors, respectively. A number of models Pannonian Basin that suggest a greater attenuation of the lithospheric have proposed the presence of a mantle diapir beneath the intra- mantle as compared to the finite extension of the crust. Horváth et al. Carpathians area (e.g., Stegena, 1967). Other models proposed rifting (1988) further improved this concept by considering radioactive heat as the driving mechanism for the Pannonian Basin (Horváth et al., generation in the crust, and the thermal blanketing effect of basin- 1975; Stegena et al., 1975), attributed to thermal thinning of the scale sedimentation. By reconstructing the subsidence and thermal Pannonian lithosphere in response to upwelling of the mantle above history, and by calculating the thermal maturation of organic matter the subducted European and Adriatic lithospheric slabs that dip in the central region of the Pannonian Basin (Great Hungarian Plain), beneath the Pannonian domain. a major step forward was made in the field of hydrocarbon prospecting At a different scale and with a different mechanical behaviour, the by means of basin analysis techniques. plume model has been also recently inferred by Burov and Cloetingh These studies highlighted difficulties met in explaining basin subsi- (2010) to be responsible for the formation of the asthenospheric up- dence and crustal thinning in terms of uniform extension, and point to- welling beneath the Pannonian basin and the initiation of continental ward the applicability of anomalous sub-crustal mantle thinning. This lithosphere subduction in neighbouring Carpathians and Dinarides issue was central to subsequent investigations involving quantitative orogens. subsidence analyses (backstripping) of an extended set of Pannonian Other models stress the back-arc position of the Pannonian domain Basin wells and cross sections and their forward modelling (Juhász with respect to the Carpathian arc and postulate that gravity-driven et al., 1999; Lankreijer et al., 1995; Lenkey, 1999; Sachsenhofer et al., passive roll-back of the subducting European lithospheric slab is the 1997). Kinematic modelling, incorporating the concept of necking driving force for the tensional subsidence of the Pannonian Basin depth and finite strength of the lithosphere during and after (e.g., Csontos, 1995; Csontos et al., 1992; Horváth, 1993; Linzer, rifting (Van Balen et al., 1999), as well as dynamic modelling studies 1996; Royden and Karner, 1984; Royden and Horváth, 1988). In a (Huismans et al., 2001b), suggested that the transition from passive modification of this model, eastward mantle flow is thought to con- to active rifting was controlled by the onset of sub-crustal flow and trol roll-back of the subducted slab and related retreat of its hinge small-scale convection in the asthenosphere. In order to quantify line. basin-scale lithospheric deformation, Lenkey (1999) carried out for- Based on thermo-mechanical finite element modelling, Huismans ward modelling applying the concept of non-uniform lithospheric et al. (2001b) were able to simulate temporal changes in rifting style stretching and taking into account the effects of lateral heat flow,

Fig. 29. Crustal thinning factors calculated by forward modelling for the Pannonian Basin applying the non-uniform stretching concept and taking the effects of lateral heat flow and flexure of the lithosphere into account (after Lenkey, 1999). Note the pronounced lateral variations in crustal extension controlling the development of deep sub-basins separated by areas of limited deformation. AM: Apuseni Mts.; DIN: Dinarides; EA: Eastern Alps; TR: Transdanubian Range; SC, WC: Southern and Western Carpathians, respectively. Local depressions of the Pannonian Basin system: Be: Bekes; Da: Danube; De: Derecske; Dr: Drava; ES: East Slovakian; Ja: Jaszsag; Ma: Mako; Sa: Sava; St: Styrian; Vi: Vienna; Za: Zala. Modified from Lenkey (1999).

flexure, and necking of the lithosphere. Calculated crustal thinning fac- final phase of basin evolution is characterized by the gradual structural tors (δ) indicate large lateral variation of crustal extension in the inversion of the Pannonian Basin system during the Late Pliocene–Qua- Pannonian Basin (Fig. 29). This is consistent with the areal pattern of ternary. During these times intraplate compressional stresses gradually the depth to the pre-Neogene basement (Horváth et al., 2006). The in- built up and caused basin-scale buckling of the Pannonian lithosphere dicated range of crustal thinning factors of 10–100% crustal extension that was associated with late- stage subsidence anomalies and differen- is in good agreement with the pre-rift palinspastic reconstruction of tial vertical motions (Horváth and Cloetingh, 1996). As evident from the Pannonian Basin, and the amount of cumulative shortening in the subsidence curves (Fig. 30), accelerated late- stage subsidence charac- Carpathian orogen (e.g., Fodor et al., 1999; Roure et al., 1993). terized the central depressions of the Little and Great Hungarian Plains As a major outcome of basin analysis studies, Royden et al. (1983) (Fig. 30b-c), whereas the peripheral Styrian and East Slovakian provided a two-stage subdivision for the evolution of the Pannonian sub-basins were uplifted by a few hundred meters after mid-Miocene Basin with a syn-rift (tectonic) phase during Early to Middle Miocene times (Fig. 30d–e) and the Zala Basin during the Pliocene–Quaternary times, and a post-rift thermal subsidence phase during the Late (Fig. 30f). The importance of tectonic stresses, both during the rifting Miocene–Pliocene. Further development of the stratigraphic database, (extension) and subsequent inversion phase (compression), is however, demonstrated the need to refine this scenario. According to highlighted by this late-stage tectonic reactivation, as well as by other Tari et al. (1999), the regional Middle Badenian unconformity, marking episodic inversion events in the Pannonian Basin (Fodor et al., 1999; the termination of the syn-rift stage, is followed by a post-rift phase that Horváth, 1995). is characterized by only minor tectonic activity. Nevertheless, the subsi- A more recent analysis of the subsidence evolution of the dence history of the Pannonian Basin can be subdivided into three Pannonian basin in the sector connecting the Carpathians with the main phases that are reflected in the subsidence curves of selected Dinarides has observed significant discrepancies with the above men- sub-basins (Fig. 30). The initial syn-rift phase is characterized by rapid tioned models by analysing the development of syn- and post-rift tectonic subsidence, starting synchronously at about 20 Ma in the sedimentation (Matenco and Radivojević, 2012). This study has entire Pannonian Basin. This phase of pronounced crustal extension is documented large-scale erosional unconformities between the classi- recorded everywhere in the basin system and was mostly limited to cally defined syn- and post-rift phases that migrate in time from the relatively narrow, fault bounded grabens or sub-basins. During the sub- Dinarides to the Carpathians, while the post-rift deposition for the sequent post-rift phase much broader areas began to subside, reflecting Early and Middle Miocene rifting is largely missing (Fig. 31). There- general down warping of the lithosphere in response to its thermal sub- fore, these authors propose that the low-angle Early–Middle Miocene sidence. This is particularly evident in the central parts of the Pannonian normal faults responsible for the basin formation are connected to a Basin, suggesting that in this area syn-rift thermal attenuation of the lower crustal detachment that was prolonged eastwards beneath lithospheric mantle played a greater role than in the marginal areas the Apuseni Mountains and affected the mantle lithosphere below (e.g., Royden and Dovenyi, 1988; Sclater et al., 1980). The third and the Transylvanian Basin. The coupling between the mechanical

Fig. 31. Depth-converted regional transects over the Pannonian Basin of Serbia and Romania. After Matenco and Radivojević (2012). extension of the Pannonian Basin and its thermal effects recorded in the system. Lower values are characteristic for the weak central part the Transylvanian Basin should have taken place in such a way that of the Pannonian Basin (5–10 km), whereas EET increases toward the the intervening Apuseni Mountains did not underwent any significant Dinarides and Alps (15–30 km) and particularly toward the Bohemian Miocene vertical movements, whose exhumation and large scale Massif and Moesian Platform (25–40 km). This trend is in good agree- dome structure is mainly the effect of Cretaceous–Paleogene deforma- ment with EET estimates obtained from flexural studies and forward tions (Merten et al., 2011). In contrast to earlier periods of extension, modelling of extensional basin formation. Systematic differences, massive thermal sag sedimentation is observed in direct relationship however, can occur and may be the consequence of significant horizon- with and post-dating the Pannonian–Lower Pontian detachments tal intraplate stresses (e.g., Cloetingh and Burov, 1996) or of mechanical and/or normal faults (such as the Makó depression, Fig. 31). The main decoupling of the upper crust and uppermost mantle that can lead to a depocenters of the thermal sag sedimentation overlie these asymmetri- considerable reduction of EET values. cal extensional structures. Furthermore, this thermal sag sedimentation The range of calculated EET values reflects the distinct mechanical has a far larger areal extent than the one of the half-grabens, covering an characteristics and response of the different domains forming part of area that corresponds to the asthenospheric upraise presently observed the Pannonian–Carpathian system to the present-day stress field. beneath the Pannonian Basin (Horváth et al., 2006). These characteristics can be mainly attributed to the memory of the This analysis demonstrated the close relationship that existed be- lithosphere. In this respect it must be kept in mind that the tectonic tween the timing and nature of stress changes in extensional basins and thermal evolution of these domains differed considerably during and structural episodes in the surrounding thrust belts, pointing to the Cretaceous–Neogene Alpine development of both the outer and mechanical coupling between the orogen and its back-arc basin. In par- intra-Carpathians units and the Neogene extension of the Pannonian allel, significant efforts have been devoted to reconstruct the spatial and Basin, resulting in a wide spectrum of lithospheric strengths. These, in temporal variations in thrusting along the Carpathian orogen (Matenco turn, exert a strong control on the complex present-day pattern of and Bertotti, 2000; Matenco et al., 2007, 2010; Roure et al., 1993; on-going tectonic activity. Schmid et al., 1998) and its relationship to foredeep depocenters (Cloetingh et al., 2005a; Matenco et al., 2003; Meulenkamp et al., 7.2.5. Deformation of the Pannonian–Carpathian system 1996; Necea et al., 2005; Tarapoanca, 2004), changes in foreland basin The present-day deformation pattern and related topography devel- geometry and lateral variations in the flexural rigidity of the foreland opment in the Pannonian–Carpathian system is characterized by lithosphere. pronounced spatial and temporal variations in the stress and strain fields (Fig. 33) (see, e.g., Cloetingh et al., 2006). Horváth and Cloetingh 7.2.4. Lithospheric strength in the Pannonian–Carpathian system (1996) established the importance of Late Pliocene and Quaternary The Pannonian–Carpathian system shows remarkable variations compressional deformation of the Pannonian Basin that explains its in crustal thickness (Fig. 22) and thermo-mechanical properties of anomalous uplift and subsidence, as well as intraplate seismicity. the lithosphere. Lithospheric rigidity varies in space and time, giving Based on the case study of the Pannonian–Carpathian system, these rise to important differences in the tectonic behaviour of different authors established a novel conceptual model for the structural reacti- parts of the system. As rheology controls the response of the litho- vation of back-arc basins within orogens. At present, the Pannonian sphere to stresses, and thus the formation and deformation of basins Basin has reached an advanced evolutionary stage as compared to and orogens, the characterization of rheological properties and their other Mediterranean back-arc basins (e.g., Bache et al., 2010)insofar temporal changes has been a major challenge to constrain and quan- as it has been partially inverted during the last few million years. Inver- tify tectonic models and scenarios. This is particularly valid for the sion of the Pannonian Basin can be related to temporal changes in the Pannonian– Carpathian region where tectonic units with a different regional stress field, from one of tension that controlled its Miocene history and rheological properties are in close contact. extensional subsidence, to one of Pliocene–Quaternary compression By conversion of strength predictions to EET values at a regional resulting in deformation, contraction, and flexure of the lithosphere scale, Lankreijer (1998) mapped the EET distribution for the entire associated with differential vertical motions. Pannonian– Carpathian system (Fig. 32). Calculated EET values are Model calculations are in good agreement with the overall topogra- largely consistent with the spatial variation of lithospheric strength of phy of the system (Fig. 33). Several flat-lying, low-elevation areas

Fig. 30. Subsidence curves for selected sub-basins of the Pannonian Basin and the Carpathian foreland. Note that after a rapid phase of general subsidence throughout the entire Pannonian Basin, the sub-basins show distinct subsidence histories from Middle Miocene times onward. Arrows indicate generalized vertical movements. O, K: Ottnangian and Karpatian, respectively (Early Miocene); BAD, SA: Badenian and Sarmatian, respectively (Middle Miocene); PAN, PO: Pannonian and Pontian, respectively (Late Miocene); PL: Pliocene; Q: Quaternary. AM: Apuseni Mts.; DIN: Dinarides; EA: Eastern Alps; PanBas: Pannonian Basin; SC, WC: Southern and Western Carpathians foreland basins, respectively. Modiﬁed from Cloetingh et al. (2006).

Fig. 32. Effective elastic thickness (EET — in km) of the lithosphere in and around the Pannonian Basin predicted from rheological calculations. BM: Bohemian Massif; MP: Moesian Platform; PB: Pannonian basin; EC, SC, WC: Eastern, Southern and Western Carpathians, respectively. Modiﬁed from Lankreijer (1998).

(e.g., Great Hungarian Plain, Sava and Drava troughs) subsided continu- Therefore, the spatial distribution of uplifting and subsiding areas ously since the Early Miocene beginning of basin development and con- within the Pannonian Basin can be interpreted as resulting from the tain 300–1000 m thick Quaternary alluvial sequences. By contrast, the build-up of intraplate compressional stresses, causing large-scale pos- periphery of this basin system, as well as the Transdanubian Range, itive and negative deﬂection of the lithosphere at various scales. This the Transylvanian Basin and the adjacent Carpathians underwent includes basin-scale positive reactivation of Miocene normal faults, late-stage uplift. and large-scale folding of the system leading to differential uplift

Fig. 33. Topography of the Pannonian–Carpathian system showing present-day maximum horizontal stress (SHmax) trajectories (after Bada et al., 2001). ‘+’ and ‘-’ symbols denote areas of Quaternary uplift and subsidence, respectively. BA: Balkanides; BM: Bohemian Massif; D: Drava Trough; DI: Dinarides; EA: Eastern Alps; EC: Eastern Carpathians; F: Foc ani Depression; MP: Moesian Platform; PB: Pannonian Basin; S: Sava Trough; SC: Southern Carpathians; TB: Transylvanian Basin; TR: Transdanubian Range; WC: Western Carpathians. Modiﬁed from Cloetingh et al. (2006).

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